Download Volatile Cycling of H2O, CO2, F, and Cl in the HIMU Mantle: A New

University of Rhode Island
DigitalCommons@URI
Graduate School of Oceanography Faculty
Publications
Graduate School of Oceanography
2014
Volatile Cycling of H2O, CO2, F, and Cl in the
HIMU Mantle: A New Window Provided by Melt
Inclusions from Oceanic Hot Spot Lavas at
Mangaia, Cook Islands
Rita A. Cabral
Matthew G. Jackson
See next page for additional authors
Follow this and additional works at: http://digitalcommons.uri.edu/gsofacpubs
Terms of Use
All rights reserved under copyright.
Citation/Publisher Attribution
Cabral, R. A., M. G. Jackson, K. T. Koga, R. F. Rose-Koga, E. H. Hauri, M. J. Whitehouse, A. A. Price, J. M. D. Day, N. Shimizu, and K.
A. Kelley (2014), Volatile cycling of H2O, CO2, F, and Cl in the HIMU mantle: A new window provided by melt inclusions from
oceanic hot spot lavas at Mangaia, Cook Island, Geochem. Geophys. Geosyst., 15, 4445-4467, doi: 10.1002/201GC005473.
Available at: http://dx.doi.org/10.1002/2014GC005473
This Article is brought to you for free and open access by the Graduate School of Oceanography at DigitalCommons@URI. It has been accepted for
inclusion in Graduate School of Oceanography Faculty Publications by an authorized administrator of DigitalCommons@URI. For more information,
please contact [email protected].
Authors
Rita A. Cabral, Matthew G. Jackson, Kenneth T. Koga, Estelle F. Rose-Koga, Erik H. Hauri, Martin J.
Whitehouse, Allison A. Price, James M.D. Day, Nobumichi Shimizu, and Katherine A. Kelley
This article is available at DigitalCommons@URI: http://digitalcommons.uri.edu/gsofacpubs/90
PUBLICATIONS
Geochemistry, Geophysics, Geosystems
RESEARCH ARTICLE
10.1002/2014GC005473
Key Points:
Volatile element concentrations have
never been reported on HIMU endmember lavas
Lavas from Mangaia represent melts
of the HIMU mantle end-member
We provide volatile and trace
element concentrations on Mangaia
melt inclusions
Supporting Information:
Readme
Tables S1–S7
Supporting information text
Filtering
Correspondence to:
R. A. Cabral,
[email protected]
Citation:
Cabral, R. A., M. G. Jackson, K. T. Koga,
E. F. Rose-Koga, E. H. Hauri,
M. J. Whitehouse, A. A. Price,
J. M. D. Day, N. Shimizu, and
K. A. Kelley (2014), Volatile cycling of
H2O, CO2, F, and Cl in the HIMU
mantle: A new window provided by
melt inclusions from oceanic hot spot
lavas at Mangaia, Cook Islands,
Geochem. Geophys. Geosyst., 15,
4445–4467, doi:10.1002/2014GC005473.
Received 25 JUN 2014
Accepted 20 OCT 2014
Accepted article online 25 OCT 2014
Published online 28 NOV 2014
Volatile cycling of H2O, CO2, F, and Cl in the HIMU mantle: A
new window provided by melt inclusions from oceanic hot
spot lavas at Mangaia, Cook Islands
Rita A. Cabral1, Matthew G. Jackson2, Kenneth T. Koga3, Estelle F. Rose-Koga3, Erik H. Hauri4,
Martin J. Whitehouse5, Allison A. Price2, James M. D. Day6, Nobumichi Shimizu7,
and Katherine A. Kelley8
1
Department of Earth and Environment, Boston University, Boston, Massachusetts, USA, 2Department of Earth Science,
University of California Santa Barbara, Santa Barbara, California, USA, 3Universit
e Blaise Pascal - CNRS - IRD, Laboratoire
Magmas et Volcans, OPGC, Clermont-Ferrand, France, 4Department of Terrestrial Magnetism, Carnegie Institution of Washington, Washington, DC, USA, 5Department of Geosciences, Swedish Museum of Natural History, Stockholm, Sweden,
6
Geosciences Research Division, Scripps Institution of Oceanography, La Jolla, California, USA, 7Department of Geology
and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA, 8Graduate School of Oceanography, University of Rhode Island, Narragansett, Rhode Island, USA
Abstract Mangaia hosts the most radiogenic Pb-isotopic compositions observed in ocean island basalts
and represents the HIMU (high m 5 238U/204Pb) mantle end-member, thought to result from recycled oceanic crust. Complete geochemical characterization of the HIMU mantle end-member has been inhibited
due to a lack of deep submarine glass samples from HIMU localities. We homogenized olivine-hosted melt
inclusions separated from Mangaia lavas and the resulting glassy inclusions made possible the first volatile
abundances to be obtained from the HIMU mantle end-member. We also report major and trace element
abundances and Pb-isotopic ratios on the inclusions, which have HIMU isotopic fingerprints. We evaluate
the samples for processes that could modify the volatile and trace element abundances postmantle melting, including diffusive Fe and H2O loss, degassing, and assimilation. H2O/Ce ratios vary from 119 to 245 in
the most pristine Mangaia inclusions; excluding an inclusion that shows evidence for assimilation, the primary magmatic H2O/Ce ratios vary up to 200, and are consistent with significant dehydration of oceanic
crust during subduction and long-term storage in the mantle. CO2 concentrations range up to 2346 ppm
CO2 in the inclusions. Relatively high CO2 in the inclusions, combined with previous observations of carbonate blebs in other Mangaia melt inclusions, highlight the importance of CO2 for the generation of the HIMU
mantle. F/Nd ratios in the inclusions (30 6 9; 2r standard deviation) are higher than the canonical ratio
observed in oceanic lavas, and Cl/K ratios (0.079 6 0.028) fall in the range of pristine mantle (0.02–0.08).
1. Introduction
Isotopic measurements performed on ocean island basalts (OIBs) have demonstrated that the Earth’s mantle
is compositionally heterogeneous [e.g., Gast et al., 1964; Allègre, 1982; Zindler et al., 1982; Zindler and Hart,
1986; Hart, 1988; Hofmann, 1997; White, 2010]. Although the cause of the chemical heterogeneity in the
mantle has been a point of discussion in the scientific community for half a century [Gast et al., 1964], it has
been suggested to arise from the subduction of oceanic and continental crust material into the mantle
[Cohen and O’Nions, 1982; Hofmann and White, 1982; White and Hofmann, 1982]. Different types of subducted material (e.g., oceanic crust, continental crust, and sediment) have been proposed to give rise to
specific mantle end-member isotopic compositions—HIMU (high m 5 238U/204Pb), EM1 (enriched mantle I),
and EM2 (enriched mantle II)—which are then melted and erupted as OIB melts at hot spots. EM2 is generally considered to be derived from continental material [White and Hofmann, 1982; Jackson et al., 2007],
but the origin of EM1 remains more controversial [e.g., Weaver, 1991; Eiler et al., 1995; Eisele et al., 2002;
Honda and Woodhead, 2005; Salters and Sachi-Kocher, 2010; Hart, 2011]. The HIMU mantle end-member is
thought to be derived from ancient subducted oceanic crust [e.g., Chase, 1981; Hofmann and White, 1982;
Zindler et al., 1982; Nakamura and Tatsumoto, 1988; Dupuy et al., 1989; Graham et al., 1992; Hauri and Hart,
1993; Hemond et al., 1994; Roy-Barman and Allegre, 1995; Woodhead, 1996; Hanyu and Kaneoka, 1997;
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4445
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Kogiso et al., 1997; Salters and White, 1998; Schiano et al., 2001; Lassiter et al., 2003; Stracke et al., 2003;
Stroncik and Haase, 2004; Nishio et al., 2005; Chan et al., 2009; Parai et al., 2009; John et al., 2010; Hanyu et al.,
2011; Kawabata et al., 2011; Salters et al., 2011; Krienitz et al., 2012; Cabral et al., 2013]. Because of the fluid
mobility of Pb during subduction zone processing, Pb loss from the slab results in high U/Pb ratios in the
subducted slab reservoir, which evolves highly radiogenic Pb-isotope compositions over time [Hofmann
and White, 1982; Zindler and Hart, 1986; Hauri and Hart, 1993; Kelley et al., 2005; Hanyu et al., 2011, 2014].
Alternatively, the high U/Pb inferred for the HIMU mantle source may be a result of U incorporation into
oceanic crust during submarine alteration [Hart and Staudigel, 1982; Hofmann and White, 1982; Chauvel
et al., 1992; Kelley et al., 2003], and this elevated U/Pb ratio preserved during subduction would evolve the
radiogenic Pb-isotopic signature associated with the HIMU mantle domain [Chase, 1981; Hofmann and
White, 1982]. A recycled origin for the HIMU mantle end-member has been further supported by sulfur isotopic measurements from olivine-hosted melt inclusions from Mangaia, which confirm that the material
erupted at Mangaia was once at Earth’s surface >2.45 billion years ago [Cabral et al., 2013].
During its life on the seafloor, oceanic lithosphere, including crust, experiences low temperature alteration
and becomes H2O-rich and CO2-rich (and gains up to several weight percent H2O and CO2 in the top portion of the crust during submarine alteration) [e.g., Staudigel et al., 1996; Alt and Teagle, 1999; Bach et al.,
2001; Wallmann, 2001; Dixon et al., 2002; Gillis and Coogan, 2011]. If the HIMU mantle end-member is truly
composed of subducted oceanic crust that is subsequently melted and erupted, this mantle end-member
can present important insights into how volatiles are returned to the mantle and recycled back to the surface at hot spots. This is important, as the distribution of volatiles in the mantle (particularly H2O) is a significant variable in determining the rheology of the mantle, the depth and extent of mantle melting, as well as
the composition of melts and how they evolve [e.g., Hirth and Kohlstedt, 1996; Gaetani and Grove, 1998; Asimow and Langmuir, 2003; Hirschmann, 2006; Portnyagin et al., 2007]. The results of Cabral et al. [2013] also
demand that the HIMU mantle sampled by Mangaia lavas experiences long-term storage in the mantle,
which has important implications for mantle dynamics and the mechanisms of long-term mantle storage
for this end-member. This discovery highlights the need for careful geochemical characterization of HIMU
lavas from Mangaia to better understand the origins of this mantle end-member. Despite its importance for
understanding mantle recycling processes, from subduction zones to hot spots, volatile measurements are
not available from the extreme HIMU mantle end-member sampled by Mangaia lavas.
Due to the lack of deep sea dredging expeditions in the Cook Islands, submarine volcanic glasses are not
available for volatile abundance characterization of HIMU mantle end-member lavas from Mangaia. Volatiles
have not been measured on HIMU mantle end-member lavas, which extend to 206Pb/204Pb values of 21.9 in
high-precision measurements made on basalts from Mangaia Island [e.g., Hauri and Hart, 1993; Woodhead,
1996; Kogiso et al., 1997; Hanyu et al., 2011]. In an effort to obtain volatile measurements of this end-member,
we have homogenized crystalline melt inclusions hosted in olivine phenocrysts extracted from subaerially
erupted whole rocks collected from Mangaia. These analyses represent the first coupled Pb isotope and major,
volatile, and trace element compositions measured on glasses from the HIMU mantle end-member.
We find that Mangaia melt inclusions host H2O and CO2 concentrations that are relatively high (1.73 wt %
and 2346 ppm, respectively; corrected for olivine fractionation). We use the new data to constrain volatile
abundances in the HIMU mantle end-member as well as volatile cycling in the mantle, from subduction
zones to hot spots.
2. Sample Descriptions, Locations, and Methods
Sample descriptions and locations, as well as a description of the geology of Mangaia, are provided in the
supporting information. A complete description of melt inclusion preparation (including the homogenization procedure) and methods for analysis of Pb-isotopes and major, trace, and volatile concentrations are
also provided in the supporting information.
3. Results
The melt inclusions in this study are alkalic (Figure 1). The forsterite values of the olivine hosting the inclusions in this study range from 75.9 to 86.1, with an average of 81.6 (supporting information Table S1). The
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4446
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
largest range in forsterite content is exhibited by olivines
from hand sample MGA-B-47
(Fo 5 80.3–86.1, supporting
information Tables S1 and S2).
The melt inclusion data presented in Figure 2 have been
corrected to be in equilibrium
with the host olivine by addition or subtraction of equilibrium olivine following the
methods outlined in Jackson
et al. [2012], and discussion of
the major, trace, and volatile
element data below refers to
abundances following this olivine fractionation correction,
unless stated otherwise (supporting information Tables S1
and S3–S5). The MgO contents
of the melt inclusions do not
span the entire range encapsulated by the whole rocks (3.10–
20.47 wt %); instead they are
Figure 1. Na2O 1 K2O versus SiO2 plotted with the alkali-tholeiite basalt line from MacDonrestricted to a range of 4.51–
ald and Katsura [1965] and applicable rock types. Mangaia melt inclusions are plotted as
10.80 wt % MgO (supporting
gray symbols and were corrected for olivine addition or subtraction (following the methods
information Table S5). The melt
outlined in Jackson et al. [2012]) to be in equilibrium with their host olivine (see supporting
information Table S5 for olivine-corrected concentrations). The whole rocks hosting the
inclusions tend to have lower
respective melt inclusions are plotted as the same symbol as the melt inclusions, but in
MgO than the host whole
black. Previously published whole rock data are shown as white circles and were obtained
rocks, and this likely owes to
from Georoc [Sarbas, 2008]. One sample (M17) from Woodhead [1996] contains anomalously
low Na2O. We argue that this is a result of significant alteration, which is supported by the
the fact that the whole rocks
high loss on ignition (LOI) of 9.33% for this sample; we have removed this sample from the
examined in this study are
figures. Whole rock data for MGA-B-25 is not available. Samples from this study plot in the
phenocryst-rich and therefore
alkalic field.
may reflect cumulate compositions. The SiO2 concentrations generally cluster in the range defined by whole rocks from Mangaia, but
some inclusions (mostly from MGA-B-47) exhibit somewhat higher SiO2 at a given MgO than the whole
rocks. Two inclusions in particular (both from hand sample MG1006) have lower SiO2 than previously identified in Mangaia lavas. The whole rock data suggest the crystallization of a Ti-rich phase in lavas where MgO
reaches <5 wt %. The melt inclusion MgO concentrations are generally above 5 wt % (with the exception of
sample MGA-B-47-RAC3 with 4.51 wt % MgO, corrected for olivine fractionation) and, with the possible
exception of MGA-B-47-RAC3, do not reach the point of extreme magmatic evolution where TiO2 concentrations are affected by fractionation of the Ti-rich phase. In sample MGA-B-47, there is one inclusion with TiO2
significantly higher than the whole rock and melt inclusion analyses. The melt inclusions generally overlap
with whole rock data in a plot of Al2O3 versus MgO (Figure 2). At a given MgO, most of the inclusions plot
within the range of whole rocks for K2O and Na2O, but a large subset of the inclusions is shifted to higher
alkali element concentrations than observed in the whole rocks. Melt inclusions generally have CaO concentrations and CaO/Al2O3 ratios that are similar to the whole rock data at a given MgO, except for inclusions
from sample MGA-B-47, which have higher CaO, and a single inclusion from MG1001, which has lower CaO.
There is a clear depletion in FeO for melt inclusions from the MGA-B-47 whole rock, which has been
observed in melt inclusions previously and predicted experimentally (discussed below) [Danyushevsky et al.,
2000; Gaetani and Watson, 2000]. Melt inclusions from the other samples lie on the FeO versus MgO trend
formed by whole rocks, except for the two melt inclusions from the MG1006 whole rock, which show an elevation in FeO compared to the rest of the melt inclusion and whole rock data. These two inclusions are
anomalous in many respects, and have the lowest SiO2, among the lowest CaO and CaO/Al2O3, and among
the highest alkali abundances in the melt inclusion suite.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4447
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Figure 2. Major element variability in Mangaia melt inclusions compared to their respective whole rocks and previously published whole
rocks from the island. Samples from this study use the same symbol convention as seen in Figure 1. Error bars in the corner of each plot
represent the average 2r measurement error of the melt inclusions. Data shown are for melt inclusions that have been corrected to be in
equilibrium with the host olivine (see supporting information Table S5 for olivine-corrected concentrations).
Previously published whole rock trace element data were obtained from the Georoc database [Sarbas,
2008] and, with few exceptions (see caption of Figure 3), were filtered for trace element analyses made by
ICP-MS (inductively coupled plasma mass spectrometry) [Hauri and Hart, 1993; Woodhead, 1996; Hanyu
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4448
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Figure 3. Trace element spidergrams of Mangaia melt inclusions compared to Mangaia whole rocks. Mangaia melt inclusion data (corrected for olivine fractionation to be in equilibrium with the host olivine, see supporting information Table S5) are plotted as black lines,
except for inclusion MG1001-5, which is plotted as a gray dashed line. All data are normalized to the chondrite-based primitive mantle values of McDonough and Sun [1995]. (a) Previously published whole rock trace element data were acquired from Georoc [Hauri and Hart,
1993; Woodhead, 1996; Sarbas, 2008; Hanyu et al., 2011] and were filtered to include only samples with a full suite of trace element data
(Ba, U, Nb, La, Ce, Pr, Nd, Sr, Zr, Hf, Sm, Y, and Yb) obtained by ICP-MS analysis (except for Zr from Hauri and Hart [1997], which was measured by XRF). A complete set of trace element data by ICP-MS did not exist from Kogiso et al. [1997], so we exclude the data from this reference from the figure. Only whole rock samples that have MgO compositions >5 wt % are plotted to avoid the effects of extreme crystal
fractionation (the majority of melt inclusions in this study have more than 5 wt % MgO). Additionally, whole rock samples with greater
than 4% LOI are also excluded from the figure to avoid effects of alteration. The estimated primary melt concentration for Mangaia, calculated from inclusion MG1001-3b following olivine fractionation corrected to be in equilibrium with mantle olivine (with forsterite 90 composition), is the following (in ppm): Ba 5 256.0, U 5 0.93, Nb 5 42.3, K 5 6511, La 5 38.0, Ce 5 68.4, Pr 5 9.3, Nd 5 33.5, Sr 5 515, Zr 5 156,
Hf 5 3.1, Sm 5 7.2, Ti 5 16,750, Y 5 28.9, Yb 5 2.3. Using the melt model outlined in section 3.2, 3% melting of the HIMU mantle source
(shown as a black dashed line) will generate this primary melt composition if the trace element concentrations in the source are the following (in ppm): Ba 5 7.7, U 5 0.028, Nb 5 1.27, K 5 195, La 5 1.15, Ce 5 2.20, Pr 5 0.35, Nd 5 1.51, Sr 5 18.1, Zr 5 12.4, Hf 5 0.29,
Sm 5 0.51, Ti 5 2520, Y 5 10.4, Yb 5 1.4. (b) Melt inclusion MG1001-5—which has an anomalous spidergram pattern—plotted as a gray
dashed line. It is compared to the average of Cape Verde calcio-carbonatites from Fuerteventura Island [Hoernle et al., 2002], which is plotted as a solid black line. Inclusion MG1001-5 has several geochemical similarities with thee calico-carbonatite average, including negative
anomalies of U, K, Zr, Hf, and positive anomalies of Nb and Sr. However, the inclusion does not have the negative Ti anomaly found in the
carbonatite.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4449
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
et al., 2011]. The whole rock trace element data and the new melt inclusion trace element data are shown
as spidergrams in Figure 3. The majority of the melt inclusion trace element data agree well with the previously published whole rock trace element data. One exception is inclusion MG1001-5 (dashed gray line in
Figure 3), which has a spidergram pattern significantly different than the other melt inclusions and previously published whole rocks. This sample was subjected to two replicate analyses by SIMS at Arizona State
University, and an additional analysis by LA-ICP-MS at URI (supporting information Table S3), and the replicate analyses gave the same spidergram pattern. A key feature of this melt inclusion is that the spidergram
has strongly negative Zr and Hf anomalies, which mimic more subtle negative Zr and Hf anomalies in HIMU
lavas from Mangaia [Kogiso et al., 1997; Hanyu et al., 2011] and other OIBs from the Cook-Austral Islands
[Chauvel et al., 1997]. Notably, this inclusion does not have anomalous major or volatile element abundances. The trace element pattern is qualitatively similar to the average calcio-carbonatite (including trace
phases, like apatite, that are associated with calcio-carbonatite) from Fuerteventura (Cape Verde hot spot)
presented by Hoernle et al. [2002], including negative anomalies U, K, Zr, Hf, and positive anomalies of Nb
and Sr. However, the inclusion does not have the negative Ti anomaly found in the calcio-carbonatites, and
we cannot offer an explanation for the origin of the anomalous trace element pattern in this inclusion.
The Mangaia melt inclusions are characterized by high H2O and CO2 contents (Figure 4a), up to 1.73 wt %
and 2346 ppm (supporting information Table S5 and Figure 4b). Inclusions from MGA-B-47 form a separate
group with lower H2O than the melt inclusions from other hand samples; melt inclusions from this sample
also appear to have lower abundances of FeO (Figure 2) and H2O. The lower H2O (and H2O/Ce, see below)
in inclusions from this sample suggest water loss from the inclusion through the host olivine (see section
4.1). An inclusion from sample MG1006 exhibits high CO2, but like MGA-B-47 it has low H2O abundances
(1750 ppm and 0.57 wt %, respectively; supporting information Table S5). However, unlike MGA-B-47 inclusions, this inclusion does not exhibit evidence for lower FeO abundance. Excluding this inclusion and all
inclusions from MGA-B-47, the majority of the Mangaia melt inclusions lie along a broad array centered at
1.5 wt % H2O (varying from 1.14 to 1.73 wt %, supporting information Table S5). This array, defined by 27
inclusions from hand sample MG1001 and three inclusions from hand sample MGA-B-25 only, spans CO2
concentrations ranging from 284 to 2346 ppm (supporting information Table S5). Excluding melt inclusion
samples from MGA-B-47 (discussed below), the melt inclusions record entrapment pressures ranging from
250 bar to approximately 1.75 kbar (see Figure 4a). This translates to entrapment depths of up to 6.4 km
(1.75 kbar) for melt inclusions with the highest concentrations of CO2 (assuming a rock density of igneous
crust of 2800 kg/m3) [Workman et al., 2006].
Figure 4c provides ratios of volatile (H2O and CO2) to nonvolatile (Ce and Nb) trace elements that have similar mineral-melt partition coefficients during mantle melting and magmatic differentiation processes.
Excluding MGA-B-47 and MG1006 inclusions, H2O/Ce varies from 119 to 245; the highest H2O/Ce value (245;
sample MG1001A12) is associated with the highest (least degassed) CO2/Nb ratio (49.5; sample
MG1001A12), but this inclusion exhibits evidence for assimilation (see below).
Koleszar et al. [2009] suggested that the MORB mantle has a CO2/Nb ratio of 300 [Koleszar et al., 2009], but
CO2/Nb in the various mantle reservoirs may not be constant, and it has also been suggested that the OIB
mantle may have higher CO2/Nb than the MORB mantle [Saal et al., 2002; Cartigny et al., 2008; Koleszar
et al., 2009]. The Mangaia melt inclusions reach a maximum CO2/Nb value of only 49.5, suggesting significant degassing (see section 4.4), and CO2 and CO2/Nb ratios in the inclusions must still be viewed as minimum values.
Fluorine concentrations range from 1073 to 1950 ppm (supporting information Table S5) in the Mangaia
melt inclusions. When compared to Nd, they are offset to higher F/Nd (30) than the F/Nd ratio of 21
observed in other MORB and OIB samples (Figure 5) [e.g., Workman et al., 2006]. F/Nd may not always reflect
mantle source characteristics if assimilation and alteration processes are involved. For example, Lassiter
et al. [2002] studied a HIMU lava from Raivavae (sample RVV318, 206Pb/204Pb 5 20.94) that showed strong
evidence for brine and oceanic crust assimilation—and melt inclusions from this sample have high F/Nd
ratios (60–183). Koleszar et al. [2009] also identified high F/Nd in five melt inclusions that range from approximately 35 to 100 and argue that this departure from the canonical ratio of fresh OIBs and MORBs of 21 may
relate to open-system behavior of F. We note that the Mangaia melt inclusion suite has elevated F/Nd ratios,
but these inclusions do not exhibit signatures associated with significant assimilation prior to entrapment
(see discussion, Figure 6).
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4450
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Cl/K2O ratios for the Mangaia melt inclusions are quite low (Figure 6). The average Cl/K2O of the melt inclusions is 0.067 (including melt inclusions from MGA-B-47), which corresponds to a Cl/K ratio of 0.079 6 0.028
and overlaps with the range for pristine oceanic lavas (Cl/K 5 0.02–0.08) found by Michael and Cornell
[1998] and Stroncik and Haase [2004].
Figure 7 compares Cl/Nb and F/Nd ratios in the Mangaia inclusions. Cl/Nb ratios, like Cl/K ratios, can be
used to monitor for assimilation of altered oceanic crust or brines (see discussion section 3.5). While the F/
Figure 4.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4451
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Nd in the Mangaia inclusions is shifted to higher values than observed in the EM1 and EM2 hot spots (Pitcairn, Samoa, and Societies), even higher F/Nd ratios are found in melts from the Azores, Galapagos, and
Australs hot spots.
Sulfides are abundant in Mangaia melt inclusions. Data for S concentrations in the silicate portions of the melt
inclusions are provided in supporting information Table S4. Since the concentration of S in the silicate portion
of the inclusions is buffered by the presence of sulfides, we cannot use the measured S concentrations to
make inferences about sulfur in the mantle source and we do not discuss sulfur further in this paper. Some
measurements of Pb isotopes, S isotopes, and major elements of the sulfides from these inclusions can be
found in Cabral et al. [2013]. Sulfides from samples MGA-B-25 and MGA-B-47 were found to have anomalous
sulfur isotopic compositions that indicate mass independent fractionation [Cabral et al., 2013].
Excluding melt inclusion measurements where the ion probe primary beam overlapped with phase boundaries (see supporting information), new measurements (supporting information Table S5) of Pb-isotopic
ratios in the Mangaia inclusions plot in a tighter cluster near the extreme HIMU mantle end-member
defined by Mangaia whole-rock compositions (Figure 8). The limited isotopic variability of the new SIMS
measurements presented here for homogenized and nonhomogenized melt inclusions is similar to that
observed in the Paul et al. [2011] LA-ICP-MS study. We do not observe the significant variability in Pbisotopic measurements in melt inclusions identified in previous SIMS studies of Mangaia inclusions [Saal
et al., 1998, 2005; Yurimoto et al., 2004]. Melt inclusions from MGA-B-25 (n 5 4 different inclusions analyzed)
show the least Pb-isotopic variability, while MGA-B-47 exhibits the most Pb-isotopic variability (n 5 12 different inclusions analyzed). MG1001 exhibits intermediate isotopic variability (n 5 23 inclusions analyzed).
Only one melt inclusion was analyzed from sample MG1002, and no melt inclusions from MG1006 were
analyzed. The sample with the greatest range in olivine forsterite contents, MGA-B-47, also has the most
heterogeneous Pb-isotopic compositions.
The new Pb-isotopic data in these melt inclusions exhibit approximately twice as much variability as
observed in high-precision Pb-isotopic measurements of leached whole rocks from the island. One melt
inclusion plots away from the other melt inclusions in Pb-isotopic space (MGA-B-47-RAC3), and expands the
melt inclusion data field significantly; without this single inclusion, the melt inclusion data field is only
slightly larger (but shifted to somewhat higher 207Pb/206Pb and 208Pb/206Pb ratios) than the data field
defined by high-precision Pb-isotopic measurements of leached whole rocks from Mangaia.
4. Discussion
In order to constrain volatile element abundances of primary magmas and to ultimately make inferences
about the composition of the HIMU mantle sampled by Mangaia lavas, it is first necessary to identify processes that operate on magmas during ascent and eruption that might modify the volatile budgets of the
magmas. In the case of melt inclusions, such processes can include degassing before entrapment of the
melt inclusion, diffusive H2O loss from the melt inclusion through the host olivine, and assimilation of
Figure 4. Volatile contents of Mangaia melt inclusions. All Mangaia melt inclusions plotted in this figure are corrected to be in equilibrium
with the host olivine (see supporting information Table S5). (a) CO2 and H2O contents of the Mangaia melt inclusions with vapor saturation
curves generated by SolEx [Witham et al., 2012] and the open-system degassing trend for alkali basalt from Dixon et al. [1997]. Major element data from one of the most primitive and undegassed samples in the suite (MG1001-5, corrected for olivine fractionation) was used
for the input values to SolEx. Melt inclusions from MGA-B-47 have experienced diffusive water loss and water data for these inclusions are
not included in subsequent figures; an inclusion from MG1006 has similarly low water, which is excluded in subsequent figures. As in
Workman et al. [2006], we interpret the nearly vertical trend (marked by steeply sloping arrow in the top plot) defined by inclusions from
sample MG1001 (n 5 27 inclusions) and MGA-B-25 (n 5 3 inclusions) to indicate CO2 degassing with negligible water loss prior to melt
inclusion entrapment (except for inclusions from MGA-B-47, which suffered diffusive H-loss post entrapment). (b) The Mangaia melt inclusion data in comparison to other OIB and MORB data sets, which consist of data obtained from submarine glasses and melt inclusions.
Data for the high CO2 inclusions from the Galapagos (Fernandina and Santiago; Koleszar et al., 2009) are ones for which CO2 from the
vapor bubble was calculated back into the melt composition, and this may explain the unusually high CO2 (extends up to 5821 ppm) in
some of the Fernandina inclusions (A. Saal, personal communication, 2014); when a similar calculation is applied to Mangaia melt inclusions, the concentration of CO2 increases by < 4% on average and < 0.02% for H2O on average (see corrected CO2 and H2O values in supporting information Table S4). Additionally, in this figure, samples were removed from plotting if the relevant data source [Clague et al.,
1995; Dixon et al., 1997; Dixon and Clague, 2001; Dixon et al., 2002; Hauri, 2002; Saal et al., 2002; Workman et al., 2006; Cartigny et al., 2008,
and references therein; Koleszar et al., 2009; Kendrick et al., 2014; Metrich et al., 2014] indicated that the sample was severely degassed due
to shallow eruption, experienced extensive assimilation, or was otherwise compromised; all samples removed, and the reason for their
removal, are outlined in the supporting information. (c) CO2/Nb and H2O/Ce ratios for Mangaia melt inclusions.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4452
Geochemistry, Geophysics, Geosystems
Figure 5. Fluorine versus neodymium of Mangaia melt inclusions in comparison to other
OIBs and MORBs. Mangaia melt inclusions are corrected to be in equilibrium with their host
olivine (supporting information Table S5) and are shown as gray symbols following the format of previous figures. Other OIB and MORB data are shown for comparison and consist of
data obtained from submarine glasses and melt inclusions [Lassiter et al., 2002; Workman
et al., 2006; Koleszar et al., 2009; Kendrick et al., 2014; Metrich et al., 2014]. Data were filtered
for degassing and assimilation as in Figure 4, with the exception of the data from the Australs: Austral samples are divided into three categories, following Lassiter et al. [2002]: Type 1
inclusions are not thought to have experienced assimilation, Type 2 inclusions (which plot
outside the figure) are thought to indicate the assimilation of Cl-rich brines, and Type 3
inclusions (which also plot outside the figure) are thought to be secondary inclusions. Mangaia melt inclusions cluster around an average value of F/Nd of approximately 30, which is
higher than the ‘‘canonical’’ F/Nd ratio for OIB and MORB of 21 [e.g., Workman et al., 2006].
MGA-B-47 data are shown because, unlike H, F does not show evidence of loss from the
inclusion.
10.1002/2014GC005473
altered oceanic crust and
brines before entrapment.
When evaluating the geochemistry of olivine-hosted
melt inclusions from Mangaia,
it is critical to evaluate all of
these processes to determine
whether primary magmatic
characteristics can be inferred
from this new suite of melt
inclusions. Identification of primary melt volatile characteristics permits inferences about
the mantle source composition
of the HIMU end-member and,
thus, volatile cycling through
the mantle resulting from
subduction-injection of oceanic crust.
4.1. Volatile Element
Concentrations: Source
Versus Process
4.1.1. Diffusive Loss of Fe
and H
Melt inclusions from whole
rock sample MGA-B-47 show
marked depletion in FeO compared to melt inclusions from
the other three whole rock
samples, which we interpret to be representative of Fe loss from the inclusions via diffusion through the
host olivine (Figure 2). Fe loss from olivine-hosted melt inclusions has been documented in other melt inclusion suites and is thought to occur following long-term storage and slow cooling in the host magma, where
the residence time is sufficient to permit Fe diffusion through the olivine [e.g., Danyushevsky et al., 1992,
2000; Sobolev and Danyushevsky, 1994; Gaetani and Watson, 2000]. The diffusivity of H in olivine is greater
than Fe, and if sufficient time elapsed for significant Fe loss to occur, then the time scale for diffusive H-loss
from the inclusions was also surpassed and H loss would have taken place provided the external melt had
lower H2O than the melt inclusions. The lower H2O in MGA-B-47 inclusions is not likely due to magma
degassing, as inclusions from MGA-B-47 have elevated CO2 (149–1335 ppm, supporting information Table
S5), so the inclusions were trapped at great depths before significant degassing of H2O. Therefore, we interpret the low H2O concentrations in MGA-B-47 inclusions (0.16–0.70 wt %, supporting information Table S5)
to be a result of diffusive water loss from the inclusions from this hand sample, and the paired Fe loss and
H2O loss in inclusions are complementary and indicative of longer time scales of storage in a magma chamber following olivine entrapment of the inclusions. Thus, in the discussion below regarding water abundances in the HIMU mantle, we exclude the inclusions from hand sample MGA-B-47.
Melt inclusions from the other hand samples exhibit no evidence for Fe loss, and have FeO abundances similar to previously published whole rocks from Mangaia. The other inclusions also have H2O (and H2O/Ce) concentrations that are higher than the H2O in MGA-B-47. Diffusive loss from the inclusions via the olivine, or
diffusive incorporation of water into the melt inclusion, is unlikely to have occurred in inclusions from MGA-B25 and MG1001. The inclusions from these two lavas form overlapping arrays in plots of CO2 versus H2O and
CO2/Nb versus H2O/Ce (Figure 4). It would seem unlikely that diffusive water loss or incorporation in inclusions
from two different lavas would generate the same arrays by coincidence, though we note that the arrays are
biased by a single inclusion with the highest CO2/Nb versus H2O/Ce. Diffusive loss of H from the inclusions
may occur following eruption during slow cooling of lavas, but it is difficult to rule out such a mechanism
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4453
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Figure 6. Mixing models demonstrating the effect of assimilating different crustal and brine materials. Mangaia melt inclusion data are plotted as gray symbols following previous figure
formatting. Loihi melt inclusion data are outlined in a light gray field and a dark gray field outlines Loihi pillow glasses [Kent et al., 1999a, and references therein]. The model lines and
end-member compositions (altered basalt, seawater, and two different brine compositions) for the assimilation model are from Kent et al. [1999a]. The Mangaia primary magma in the
model uses olivine fractionation corrected data from melt inclusion MG1001-14 to define the primary basalt composition because it has the most primitive MgO content (10.8 wt %, corrected for olivine fractionation, see supporting information Table S5) in the inclusion suite. Like most of the inclusions in this study, MG1001-14 has a relatively low Cl/K2O ratio (0.064)
typical of mantle-derived lavas that have not experienced significant assimilation of hydrothermally altered material. Notably, the composition of one melt inclusion (MG1001A12, which
also has the highest H2O/K2O and H2O/Ce in the melt inclusion suite) may be explained by the assimilation of altered basalt and/or seawater, and this sample is indicated in all plots of
this figure. Data from MGA-B-47 and MG1006 inclusions are not shown owing to diffusive water loss.
lowering the H2O in the melt inclusions from the primary magmatic values. For example, there is no clear relationship between H2O and Ce in the inclusions. However, owing to the fact that small melt inclusions are
more susceptible to diffusive loss of H2O than larger inclusions [Chen et al., 2011], a positive relationship
between H2O and melt inclusion diameter will be present if diffusive loss of H2O has occurred. We observe no
such relationship in inclusions from MGA-B-25 and MG1001 (Figure 9), which argues against diffusive loss of
H2O from the inclusions; however, the inclusions from sample MGA-B-47 that have suffered Fe loss and H2O
loss do show a weak relationship between melt H2O and melt inclusion diameter, which is consistent with diffusive loss of H from inclusions in this sample. However, while diffusive loss from MGA-B-25 and MG1001
inclusions seems unlikely (given the lack of relationship between H2O and inclusion size), it cannot be ruled
out, and the H2O/Ce ratios in the melt inclusions should be treated as minima. For inclusions from samples
MGA-B-25 and MG1001, we also conclude that diffusive exchange through the olivine did not modify the
abundance of the other incompatible trace elements, like F, Cl, and S, because their diffusivities are expected
to be significantly slower than those of Fe and H [Cherniak, 2010; Bucholz et al., 2013].
After excluding inclusions from MGA-B-47, H2O/Ce ratios exhibit some variability in Mangaia inclusions.
While H2O concentrations and H2O/Ce ratios are relatively high in the Mangaia melt inclusions (excluding
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4454
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
inclusions from sample MGA-B47), some H2O loss may occur
by diffusive H2O loss from
inclusions [e.g., Chen et al.,
2011; Gaetani et al., 2012] and
slight loss of H2O can occur via
degassing prior to inclusion
entrapment (see below), and
this may explain the lower
H2O/Ce ratios (down to 119)
observed in some of the Mangaia inclusions.
Figure 7. Cl/Nb versus F/Nd for Mangaia melt inclusions compared to global OIBs and
MORBs. Mangaia melt inclusion data are plotted as gray symbols following previous figure
formatting. The dashed box represents the canonical F/Nd value from Workman et al. [2006]
and the range of Cl/Nb suggested for the mantle [Lassiter et al., 2002; Saal et al., 2002; Workman et al., 2006; Koleszar et al., 2009; Shaw et al., 2010; Kendrick et al., 2014; Metrich et al.,
2014, www.earthchem.org/petdb]. Data from other locations are provided for comparison
and consist of data collected from submarine glasses and melt inclusions [Lassiter et al.,
2002; Workman et al., 2006; Koleszar et al., 2009; Kendrick et al., 2014; Metrich et al., 2014].
See the caption of Figure 5 for description of Austral sample groups. Data are filtered in the
same way as in the caption of Figure 4.
Figure 8. 208Pb/206Pb versus 207Pb/206Pb measurements on Mangaia melt inclusions from
this study. Data from this study obtained on homogenized inclusions are shown as solid
black circles (supporting information Table S5). Data from sulfide inclusions are plotted as
white diamonds; the average of two analyses is plotted for one of the sulfide analyses (supporting information Table S5). FOZO values are estimated and are from Hart et al. [1992].
The field for previously published whole rock basalt data from Mangaia shows only the data
obtained from leached whole rocks that were measured by MC-ICP-MS [Hanyu et al., 2011]
or by a double-spike TIMS method [Woodhead, 1996]. The error ellipse (in the bottom right
portion of the figure) represents the 2r reproducibility (but not the absolute value) of the
secondary standard measurements during the session (these measurements are shown individually in supporting information Figure S2).
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4.1.2. Degassing of H2O
and CO2
Models examining solubility of
H2O and CO2 in mantle melts
have shown that, during opensystem degassing, the vapor
phase is dominated by CO2
until very low pressures (100
bars) [Dixon, 1997]. As in Dixon
[1997] and Workman et al.
[2006], we interpret the nearly
vertical degassing trend (Figure 4a) and entrapment pressures >100 bar to indicate
negligible H2O loss from the
melt inclusions from hand samples MG1001 and MGA-B-25.
Owing to the low solubility of
CO2 in mantle melts, in only
rare cases has CO2 been
observed to be undersaturated, thereby providing information about primary melt
CO2 abundances [Saal et al.,
2002]. In the case of Mangaia,
and all other OIB and MORB
locations examined to date,
extensive degassing of CO2 is
evident. Therefore, primary
melt CO2 abundances cannot
be calculated with the present
data set, as estimates determined for the primary melt
CO2 concentration using CO2/
Nb ratios have large uncertainties owing to variability in the
CO2/Nb ratio [Saal et al., 2002;
Cartigny et al., 2008; see section 4.4), particularly if carbonatite is present during
melting [Dasgupta et al., 2009].
4455
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Assimilation during magma
ascent: Assimilation of oceanic
crust and/or brines during
ascent and eruption of magma
can modify the Cl and H2O
abundances of the magma
[Michael and Schilling, 1989;
Jambon et al., 1995; Michael
and Cornell, 1998; Kent et al.,
1999a, 1999b; Lassiter et al.,
2002; Stroncik and Haase, 2004;
Kendrick et al., 2013a]. Kent
et al. [1999a] explored this
issue through the analysis of
matrix glass and olivine-hosted
glass inclusions from Loihi seamount, Hawaii. Following Kent
et al. [1999a], we examined the
Figure 9. Melt inclusion H2O contents compared to melt inclusion long axis diameter. Manvariability of Cl/K2O and H2O/
gaia melt inclusion data are plotted as gray symbols following previous figure formatting.
K2O in the melt inclusions from
The data have been corrected to be in equilibrium with the host olivine. The lack of correlation for melt inclusions from samples MGA-B-25 and MG1001 suggest no significant diffuMangaia (Figure 6). The majorsive H2O loss from the inclusion. However, the negative correlation observed for melt
ity of melt inclusions plot in a
inclusions from hand sample MGA-B-47 does suggest diffusive H2O loss, and is consistent
small cloud near the most
with the previous observations from this study.
primitive (highest MgO) Mangaia melt inclusion composition (see Figure 6 caption for description) and generally exhibit limited variability, especially in comparison with the variability seen in the Loihi samples, which are known to have
experienced assimilation. However, we note that assimilation of altered basalt, as shown in Figure 6, can
increase the H2O/K2O (and H2O/Ce) without significantly increasing the Cl/K2O, and thus Cl/K2O alone cannot be used to evaluate assimilation of altered oceanic basalt. However, in plots showing B/K2O, H2O/K2O,
B/Cl, and H2O/Cl, one inclusion (MG100A12) is systematically shifted in a direction that suggests assimilation
of altered oceanic crust and seawater (Figure 6). This inclusion also has the highest H2O/Ce (245) in the
suite, and the high value may owe to assimilation of H2O. However, the other inclusions do not appear to
be shifted consistently in the direction of assimilation of crustal materials, brines, or seawater. These inclusions, which exhibit no clear evidence for assimilation, have H2O/Ce values that range up to 200. However,
we cannot exclude the possibility that minor assimilation has occurred in other inclusions in the Mangaia suite,
which may result in the incorporation of non-magmatic H2O. If this is the case, then the H2O/Ce of the Mangaia HIMU melts will be overestimated. Further work analyzing HIMU samples at other localities, particularly
quenched submarine glasses, will be important for constraining the H2O content of this mantle end-member.
4.2. H2O in the HIMU Source and Constraints on Cycling of Subducted Oceanic Crust
The crustal recycling origin hypothesis for the HIMU mantle end-member requires that material (oceanic
crust) once at the Earth’s surface was processed through a subduction zone, experienced long-term storage
in the mantle, and was reerupted at an OIB locality, such as Mangaia. The input of water-rich altered oceanic
crust to subduction zones may represent a significant contribution of H2O to the mantle [e.g., Hacker, 2008],
and data from the Mangaia melt inclusions can provide constraints on the water that remains in the oceanic
crust during subduction.
In order to calculate representative H2O abundances for the HIMU mantle source, it is important to consider
the melting process itself. Owing to the incompatibility of H2O, mantle melting concentrates water in the melt
relative to the mantle source and subsequent fractional crystallization further increases the concentration of
H2O in the melt. Comparing H2O concentrations to the concentrations of other similarly incompatible elements (e.g., Ce and K) provides a means to ‘‘see’’ through the melting process, as H2O/Ce ratios are not significantly modified during mantle melting to generate OIB and MORB [e.g., Michael, 1995; Dixon et al., 2002].
When H2O is plotted against K2O and Ce, lavas that erupt at different hot spot localities fall on different
trends (Figure 10). Melt inclusions from Mangaia have relatively high H2O concentrations compared to other
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4456
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
submarine glasses and melt
inclusions in the global OIB
and MORB data set, but have
intermediate Ce and K2O concentrations; lavas erupted at
EM1 and EM2 localities also
tend to have elevated H2O, but
have higher Ce and K2O at a
given H2O abundance. Lavas
from the enriched mantle hot
spots, which host elevated
87
Sr/86Sr and lower 143Nd/144Nd
ratios [Zindler and Hart, 1986;
Hofmann, 1997], exhibit Ce
concentrations that are nearly 2
times higher at a given H2O
abundance than the HIMU
melts.
Plots of H2O/Ce versus Sr and
Pb isotopes also show fields for
a number of OIB and MORB
localities (Figure 11). Mangaia
lavas and melt inclusions
anchor the low 207Pb/206Pb
and low 208Pb/206Pb region of
the global data set, and in situ
SIMS measurements of Pb isotopes on the Mangaia melt
inclusions from this study have
relatively homogeneous compositions that are similar to
whole rocks from the island
[Nakamura and Tatsumoto,
1988; Dupuy et al., 1989; Hauri
and Hart, 1993; Woodhead,
Figure 10. H2O versus Ce and K2O compared to global OIB and MORB data. Mangaia melt
inclusion data are plotted as gray symbols following previous figure formatting, and have
1996; Kogiso et al., 1997; Tatbeen corrected to be in equilibrium with the host olivine (supporting information Table S5).
sumi et al., 2000; Schiano et al.,
The H2O contents found in Mangaia melt inclusions are among the highest for OIBs and
2001; Hanyu et al., 2011]. ESC
MORBs. Data arrays from other OIB and MORB localities are plotted for comparison and consist of data collected from submarine glasses and melt inclusions [Schilling et al., 1985; Clague
lavas with the most radiogenic
et al., 1995; Douglass et al., 1995; Dixon et al., 1997; Langmuir et al., 1997; Pan and Batiza, 1998;
Pb-isotopic compositions
Dixon and Clague, 2001; Dixon et al., 2002; Hauri, 2002; Kingsley, 2002; Saal et al., 2002; Simons
et al., 2002; Yang et al., 2003; le Roux et al., 2002; Workman et al., 2004, 2006; Cartigny et al.,
(206Pb/204Pb, 207Pb/204Pb, and
2008, and references therein; Koleszar et al., 2009; Kelley et al., 2013; Kendrick et al., 2014;
208
Pb/204Pb up to 19.91, 15.64,
Metrich et al., 2014]. Easter microplate (EMP) and Easter seamount chain (ESC) are abbreviated
and 39.66, respectively, which
for clarity. Kilauea glass and Hawaii melt inclusion samples [Clague et al., 1995; Hauri, 2002] do
not have trace element data and therefore do not plot in the H2O versus Ce plot. The volatiles
translates to 208Pb/206Pb 5 1.99
database was filtered in the same way as described in the caption of Figure 4.
and 207Pb/206Pb 5 0.785) have
H2O/Ce ratios (up to 244) that are similar to the highest H2O/Ce ratios (200) in the subset of Mangaia melt
inclusions that show no evidence for assimilation. We note that other localities have even higher H2O/Ce,
including Pico Island in the Azores (H2O/Ce 5 340) [Metrich et al., 2014] and N. Atlantic MORB (350) [Dixon
et al., 2002], which indicates that the HIMU mantle reservoir does not have the highest H2O/Ce.
We note that Dixon et al. [2002] argued for a lower H2O/Ce ratio for the HIMU mantle of 100, but this is a factor of two lower than the highest H2O/Ce values from the ESC lavas with the most radiogenic Pb-isotopic
compositions [Dixon et al., 2002] and the Mangaia inclusions (200). In contrast to the HIMU mantle, H2O/
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4457
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Ce values for EM-type (high 87Sr/86Sr) lavas from Societies, Pitcairn, and Samoa extend to the low H2O/Ce
ratios (H2O/Ce < 100) [Dixon et al., 2002; Workman et al., 2006; Kendrick et al., 2014]. Thus, the EM mantle signature anchors the high 87Sr/86Sr, low H2O/Ce portion of the global mantle array. Sr-isotopes were not
measured on the Mangaia inclusions, but if the Sr-isotopic compositions of the inclusions are similar to
whole rocks from the island (just as Pb-isotopic compositions of the inclusion suite examined here are similar to the Pb-isotopic compositions of the whole rocks from Mangaia), then Mangaia, together with the ESC
lavas with the most radiogenic Pb-isotopic ratios, anchors the low 87Sr/86Sr and high H2O/Ce portion of the
global mantle array.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4458
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Given a H2O/Ce ratio of the HIMU mantle of 200, it is possible to calculate the H2O concentration of this
mantle end-member if a reliable estimate for the mantle source Ce can be obtained. There are some
bounds that can be placed on the Ce concentration of the HIMU mantle that can be used to constrain the
H2O abundance in the HIMU mantle source. It is important to note that the Ce concentrations for the HIMU
mantle source are not known, and there are many uncertainties associated with estimating the abundances
of incompatible elements in the mantle source of lavas. The following discussion uses estimates from the literature, and assumes that the HIMU mantle is a mixture of altered oceanic crust [Hoffman and White, 1982]
and a depleted plume component, which is similar to earlier models for this mantle source [Hauri and Hart,
1993; Stracke et al., 2003]. It is reasonable to assume that the Ce concentration will not be more depleted
than the depleted plume component, which is estimated to have a Ce concentration of 1.3 ppm [Jackson
and Jellinek, 2013], but not more enriched than altered oceanic crust, which is estimated to have a Ce concentration of 6 ppm (as provided by Dixon et al. [2002]). The proportion of the two end-members (altered
oceanic crust and depleted plume component) contributing to the HIMU mantle source is estimated from
prior work suggesting a two-component mixture for generating the HIMU Mangaia mantle source [e.g.,
Hauri and Hart, 1993]. If we assume the Mangaia HIMU mantle has a 20% contribution of recycled oceanic
crust, which is consistent with coupled Os-isotopic and Pb-isotopic constraints on Mangaia lavas [Hauri and
Hart, 1993], and if the remaining 80% of the Mangaia source is composed of the depleted plume component, then the Mangaia HIMU source has a Ce concentration of 2.2 ppm. Isotopic constraints suggest that
higher proportions of recycled oceanic crust in the HIMU source are unlikely, particularly if the Mangaia
mantle source is >2.5 Ga [Hauri and Hart, 1993; Day et al., 2009, 2010]. Dixon et al. [2002] assumed a Ce concentration of 6 ppm for the HIMU source, which assumes that HIMU is pure recycled altered oceanic crust
without a depleted mantle contribution; other estimates for the Ce concentration of altered oceanic crust
and MORB are available in the literature [Kelley et al., 2003; Gale et al., 2013], but we use the estimate from
Dixon et al. [2002] for direct comparison with their model results exploring the H2O abundances in the mantle. We emphasize that the bulk major element composition of HIMU lavas, and the compositions of the olivines, are not well matched by melting a source comprised of purely mafic material [Jackson and Dasgupta,
2008; Herzberg et al., 2014], and we do not advocate a model for the melting of a purely mafic lithology in
the Mangaia mantle; instead, Herzberg et al. [2014] argue that the HIMU mantle consists of fertile peridotite
enriched by mixing completely with a minor component of pyroxenite. Similarly, the contribution of 20%
mafic material to the HIMU source is assumed to be mixed with the depleted-mantle peridotite such that
the final mixture is an enriched peridotite [Herzberg, 2011]. Using the estimated H2O/Ce ratio for Mangaia
primary melts (200), and a Ce concentration of 2.2 ppm, the H2O concentration in the HIMU mantle source
is estimated to be 440 ppm.
The primary uncertainty in the estimate of H2O concentration of the HIMU source derives from uncertainty
in the source Ce concentration, which in our calculation scales with the amount of recycled oceanic crust in
the HIMU mantle. We have adopted a recycled oceanic crust mass fraction of 20% in our calculation based
on isotopic constraints: a higher proportion of recycled oceanic crust will yield a higher estimated Ce concentration and a higher HIMU source H2O concentration, and a lower proportion of oceanic crust yields a
lower HIMU source H2O concentration.
Using a simple melt model, a HIMU mantle source water content of 440 ppm can easily generate the H2O
abundances estimated for Mangaia primary melts. Following olivine fractionation correction so that the
melt is in equilibrium with mantle olivine (Fo90), the inclusion with the highest H2O abundance (MG1001-3b
Figure 11. H2O/Ce versus 87Sr/86Sr, 207Pb/206Pb, and 208Pb/206Pb isotopes for Mangaia melt inclusions and global OIBs and MORBs. The Pbisotopic data for the Mangaia inclusions (supporting information Table S5) are relatively homogeneous and similar to Mangaia whole rock
values. Therefore, while Sr-isotopes were not measured in the inclusions, they are also assumed to be homogeneous and similar to the Mangaia whole rock values. (a) We applied an average 87Sr/86Sr to all of the melt inclusions using data from Hanyu et al. [2011] for age corrected,
leached whole rocks. Data from other OIB and MORB sources are provided for comparison and consist of submarine glass samples and melt
inclusions [Staudigel et al., 1984; Hannan and Schilling, 1989; Devey et al., 1990; Dosso et al., 1991; Fontignie and Schilling, 1991; Woodhead and
Devey, 1993; Douglass et al., 1995; Dixon et al., 1997; Kingsley and Schilling, 1998; Douglass et al., 1999; Dosso et al., 1999; Frey et al., 2000; Dixon
and Clague, 2001; Dixon et al., 2002; Kingsley, 2002; Simons et al., 2002; Yang et al., 2003; Honda and Woodhead, 2005; Workman et al., 2006;
Cartigny et al., 2008, and references therein; Hanyu et al., 2011; Kelley et al., 2013; Kendrick et al., 2013a, 2014]. Not all data from Figure 10 are
shown in this figure because isotopic compositions are not available. The box labeled ‘‘HIMU’’ is the estimated range of the HIMU source
value, and excludes the Mangaia inclusion with the highest H2O/Ce (245; inclusion MG1001A12) owing to possible assimilation of seawater
or altered oceanic crust. The volatiles database was filtered in the same way as described in the caption of Figure 4.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4459
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
has 1.73 wt % H2O, when corrected to be in equilibrium with the host olivine; supporting information Table
S5) has a water concentration of 1.37 wt % (and an H2O/Ce ratio of 200). Using mineral-melt partition
coefficients and a mantle source mineralogy with clinopyroxene (12% modal abundance), orthopyroxene
(25%), olivine (55%), and garnet (8%) modes provided in Kelemen et al. [2004] and summarized in Jackson
et al. [2010], and assuming that H2O and Ce have the same mineral-melt partition coefficients during mantle
melting, we use a modal aggregated fractional melting model to explore the conditions under which a
mantle source with 440 ppm H2O can generate melts with 1.37 wt % H2O.
If the HIMU source has 440 ppm H2O, then the water abundance of the HIMU primary melt (1.37 wt %) is
achieved when the melt fraction is 3.0%. This degree of melting is within the range previously suggested
for melt generation in the Cook-Australs [Herzberg and Gazel, 2009].
The melt model can also be used to provide an estimate of the refractory lithophile element concentrations
in the Mangaia mantle source. When the melt fraction is 3%, the melt model yields a Ce concentration of
68.5 ppm in the melt (assuming 2.2 ppm Ce in the HIMU source). This value is identical to the Ce concentration (68.5 ppm) for the inclusion with the highest H2O (MG1001-3b) following olivine fractionation correction so that the melt is in equilibrium with mantle olivine (Fo90), and this trace element composition is
adopted as the Mangaia primary melt composition (see Figure 3 caption). Because the melt model does a
reasonable job of matching the Ce concentration in Mangaia primary melts, we use the same melt model
(with 3% melting), together with the estimated Mangaia primary melt trace element composition, to
invert for the source composition of the other trace elements in the primary melt shown in Figure 3 (and
presented in the caption). Using the specified melt model, this source will generate the trace element abundances in the Mangaia primary melt composition when the melt fraction is 3%. This source composition is
presented in the caption of Figure 3.
4.3. Dehydration of the Subducted Slab to Generate the HIMU Source
Altered oceanic crust is estimated to have a H2O/Ce ratio of 2500–5000 [Dixon et al., 2002] and a water concentration of 2–3 wt %. If the HIMU mantle hosts a significant component of recycled oceanic crust, then an
important question remains: How did the H2O/Ce ratio become so low in the HIMU mantle? Subducted
material experiences a two-stage dehydration process: (1) stage one is defined by subduction zone processing [e.g., Hacker, 2008; van Keken et al., 2011], where H2O is lost from the slab, and (2) stage two is defined
by diffusion of water out of the subducted slab during long-term storage in the mantle [e.g., Workman
et al., 2006]. Following Dixon et al. [2002], we can look at the net effect of these two processes by comparing
presubduction H2O/Ce ratios of mature oceanic crust to the H2O/Ce ratios in HIMU melt inclusions erupted
at Mangaia. Compared to the presubduction mature oceanic crust H2O/Ce of 2500–5000, we find that the
HIMU mantle source today, which is suggested to be a mixture of recycled material and depleted plume
component mantle [e.g., Hauri and Hart, 1993], has an H2O/Ce ratio of 200 (which is a factor of two higher
than the H2O/Ce ratio suggested for the HIMU source [100] by Dixon et al. [2002]). A reduction of the H2O/
Ce ratio from 2500–5000 to 200 represents a factor of 12.5–25 reduction in original H2O content in the subducted slab compared to the loss of Ce.
It is useful to deconvolve the relative effects of stage one of dehydration (subduction zone processing) from
the effects of stage two of dehydration (long-term mantle storage, diffusive water loss). van Keken et al.
[2011] reported estimates of water content entering subduction zones and surviving to reach >230 km
depth based on petrological models and global subduction zone rock fluxes. Using H2O values of igneous
crust from before subduction (9.3 3 108 Tg/Myr) [Hacker, 2008] and after subduction (1.83108 Tg/Myr) [van
Keken, 2011], igneous crust is estimated to dehydrate 81% from the subduction process alone. This constrains long-term storage dehydration of the slab by diffusive water loss to account for the remaining reduction of water in the HIMU mantle reservoir, which can efficiently reduce H2O contents in subducted slabs
on geologic time scales (see below).
Diffusion rates for trace elements (like Ce) are very low and heterogeneities in the solid mantle will not equilibrate with these elements on length scales of more than 10 m over the age of the Earth [e.g., Hofmann
and Hart, 1978; Van Orman et al., 2001]. In contrast, hydrogen diffuses quickly and the H2O/Ce ratio can be
reduced by diffusive loss of H2O into the ambient mantle. Using the H2O concentration values of altered
oceanic crust from Dixon et al. [2002] (2–3 wt %) and the amount of dehydration calculated above (81%),
the recycled slab after the subduction zone would have 0.38–0.57 wt % (3800–5700 ppm) H2O. If the
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4460
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Mangaia mantle, which is estimated to have 440 ppm water, is composed of recycled slab (20% by mass)
and a depleted component (80%) that has a water content similar to ambient mantle (116 ppm) [Salters
and Stracke, 2004], then the slab component in the Mangaia mantle has a water content of 1740 ppm.
This is 2.2–3.3 times lower than the estimated water content of a slab subducted to postarc depths, and we
argue that diffusive loss from the slab during long-term storage in the mantle can reduce the water content
in the slab from postsubduction values.
We model the effect of diffusion in a manner similar to Workman et al. [2006]. A diffusion calculation [see
Zhang, 2008, equation (3-48b)] was performed assuming the placement of postsubduction mature oceanic
crust in an infinite reservoir of depleted mantle. We estimate depleted mantle to have 116 ppm H2O [Salters
and Stracke, 2004], adopt a thickness of 7 km for the slab, and assume a storage time of 2.5 Ga. The range of
water contents remaining in the slab after subduction (0.38–0.57 wt % H2O) are used as inputs for the diffusion calculation. The diffusion coefficient of H in olivine was found to be 1.4 3 1028 m2/s at 1600 C for the
[100] (fast) axis by Mackwell and Kohlstedt [1990]. In this simple diffusion calculation, following subduction,
the slab experiences approximately 90% diffusive loss of water during 2.5 Ga of residence in the mantle: the
water content of the slab is reduced from 3800 to 5700 ppm to approximately 340 to 450 ppm. Thus, diffusive loss of H2O from the slab following subduction zone dehydration would make the slab much drier
(340–450 ppm H2O) than our model for the slab (1740 ppm H2O) in the HIMU mantle. However, if the slab
is thicker, or slabs pile up on each other in the mantle (effectively thickening the diffusive distance), then
the diffusive loss of H2O would be reduced. Additionally, we note that pyroxene-rich and garnet-rich lithologies have a higher water storage capacity than olivine-rich lithologies [Hirschmann et al., 2009; O’Leary et al.,
2010; Ardia et al., 2012], and higher modal pyroxene and garnet abundances in the HIMU mantle source
may allow the preservation of higher water abundances during long-term storage in the mantle. This is not
an unreasonable model, as HIMU sources have long been suggested to host a significant mafic component
[Hofmann and White, 1982; Hauri and Hart, 1993; Day et al., 2009] or to be a fertile peridotite that has been
enriched by mixing completely with a component of pyroxenite (thereby increasing the pyroxene abundance in the enriched peridotite resulting from the mixture) [Herzberg et al., 2014].
4.4. CO2 in the Generation of HIMU Lavas
It is apparent that the CO2 concentrations in Mangaia melt inclusions are high compared to CO2 abundances measured in many other OIB settings (Figure 4b). In fact, carbonate blebs were previously noted in crystalline melt inclusions hosted in olivines [Saal et al., 1998; Yurimoto et al., 2004], and Saal et al. [1998] found
that melt inclusions from two of the whole rocks (MGA-B-47 and MGA-B-25) examined in this study host carbonatite blebs (unfortunately, we did not examine these particular melt inclusions). We infer that some
inclusions in the present study may have also hosted carbonate blebs prior to homogenization, but we did
not observe carbonatite blebs in the inclusions posthomogenization. Saal et al. [1998] suggested that carbonatite plays a key role in the genesis of HIMU lavas, and Hauri et al. [1993] found that carbonatite plays an
important role in modifying mantle xenoliths from the island of Tubuai (located in the Cook-Australs). The
association of carbonate with lavas derived from the HIMU mantle source could be an important clue for
ultimately understanding the origin of this mantle end-member [Jackson and Dasgupta, 2008].
Partial melt compositions of carbonated mafic lithologies are shown experimentally to give rise to melts
that are similar to the inferred compositions of HIMU primary melts (matching geochemical characteristics
such as SiO2, TiO2, Al2O3, CaO, Na2O, and CaO/Al2O3) [Dasgupta et al., 2004; Gerbode and Dasgupta, 2010;
Mallik and Dasgupta, 2012]. It has also been suggested that there is interaction of this melt product with surrounding peridotite, which leads to a closer match for FeO and is consistent with a recent study of the
HIMU mantle source lithology by Herzberg et al. [2014] [Dasgupta et al., 2004; Gerbode and Dasgupta, 2010;
Mallik and Dasgupta, 2012]. Experimental evidence shows that it is difficult to generate strongly SiO2-undersaturated and high CaO/Al2O3 melt compositions, like those observed in HIMU lavas, without high concentrations of CO2 in the mantle source [Kushiro, 1975; Eggler, 1978; Dasgupta et al., 2004, 2007; Gerbode and
Dasgupta, 2010; Mallik and Dasgupta, 2012], which strongly suggests that the HIMU mantle end-member is
CO2-rich.
Unfortunately, it is not possible to reliably estimate the CO2 abundance in the primary melts of the HIMU
mantle owing to large uncertainties in the CO2/Nb ratio that has been used to estimate mantle source CO2
abundances [Saal et al., 2002]. The CO2/Nb ratio of the depleted mantle is estimated to be 300 [Koleszar
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4461
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
et al., 2009], which is slightly higher than the value of 239 that was originally proposed [Saal et al., 2002],
and this is likely a lower limit for the CO2/Nb ratio of the OIB mantle. Cartigny et al. [2008] observed a relationship between CO2/Nb and (La/Sm)N in MORB. If we assume that this relationship applies to the Mangaia
melts, which have an average (La/Sm)N of 2.9 (supporting information Table S5), then the CO2/Nb of the
Mangaia mantle source is 1000; this value, when paired with the estimated Nb concentration of the HIMU
mantle source (see caption of Figure 3), yields unreasonably high CO2 in Mangaia primary melts. If carbonatite is present during melting, CO2 and Nb will have dramatically different partition coefficients, and CO2/Nb
ratios in the primary melt will exceed 104 [Dasgupta et al., 2009]. This serves to underscore the great uncertainty in the CO2/Nb ratio of primary melts, particularly when carbonatite is present.
While it is not possible to reliably estimate the CO2 in abundance in the HIMU mantle source or in HIMU primary melts, the HIMU mantle is thought to have elevated CO2 [Hauri et al., 1993; Saal et al., 1998; Jackson
and Dasgupta, 2008; Castillo, 2014]. The origin of this elevated CO2 may be traced to the recycled crustal origin of the HIMU mantle, where CO2-rich and H2O-rich altered oceanic crust [e.g., Staudigel et al., 1996; Alt
and Teagle, 1999; Bach et al., 2001; Wallmann, 2001; Dixon et al., 2002; Gillis and Coogan, 2011] is subducted
into the mantle and recycled into the origin of HIMU lavas. The carbonated nature of the HIMU mantle
source is consistent with experimental work attempting to model the major element composition of HIMU
lavas [Dasgupta et al., 2004; Gerbode and Dasgupta, 2010; Mallik and Dasgupta, 2012] and has important
implications for the deep storage of recycled carbon in the mantle.
4.5. F/Nd and Cl/Nb Signature of HIMU Melts
The F and Cl concentrations in melt inclusions are useful for constraining the undegassed, preeruptive compositions because their degassing pressure is typically less than 100 bars [Spilliaert et al., 2006]. Additionally,
F and Cl are not thought to diffuse in or out of the host olivine [Bucholz et al., 2013]. Normalizing F and Cl to
elements with similar bulk partition coefficients, namely Nd and Nb, respectively [Dalou, 2011], the F/Nd
and Cl/Nb ratios measured in melt inclusions have been used to characterize those of the source [e.g., RoseKoga et al., 2012]. Known as a canonical ratio, F/Nd is largely insensitive to mantle melting, and natural samples (melt inclusions and glasses from MORB and OIB) display a relatively constant value of 21 over a range
of Nd concentrations that vary over two orders of magnitude [Workman et al., 2006]. In cases where a melt
inclusion has undergone assimilation of crustal materials or brines, F/Nd as high as 60–183 have been
observed [Lassiter et al., 2002]. However, such inclusions are often accompanied by extremely elevated Cl/
Nb ratios (from 1271 to 7166; see section 2) [Lassiter et al., 2002]. Typically, Cl/Nb ranges from 5 to 35 among
the data of MORB and OIB glasses and melt inclusions [Lassiter et al., 2002; Saal et al., 2002; Workman et al.,
2006; Koleszar et al., 2009; Shaw et al., 2010; Kendrick et al., 2014; Metrich et al., 2014; www.earthchem.org/
petdb]. Notably, once altered oceanic crust assimilation has been identified and excluded, any deviation
from the relatively small range of these geochemical ratios indicates the presence of material that is not
considered to be typical mantle.
Mangaia melt inclusions display Cl/Nb values similar to MORB melt inclusions and samples from other OIBs,
such as the Azores and Samoa (Figure 7). The Cl/Nb ratio of the source of Mangaia is likely similar to that of
the MORB source. On the contrary, Mangaia melt inclusions show a higher F/Nd (30 6 9; 2r standard deviation) than the ‘‘MORB-OIB’’ canonical value of 21, and the Mangaia inclusions range up to a F/Nd ratio of 41.
The higher F/Nd values are mainly related to higher F concentrations in the Mangaia melt inclusions (about
30% higher), while Nd concentrations are comparable to other OIB data (e.g., 45 6 5 ppm; the Azores, Samoa,
and this study). The Nd concentration in the HIMU source is estimated to be 1.38 ppm (see Figure 3), which,
when combined with the average Mangaia melt inclusion F/Nd ratio (30), yields F abundances in the HIMU
source of 41 ppm. By comparison, the F concentrations of DMM and chondrite-based primitive mantle are
estimated to be 11 and 25 ppm, respectively [McDonough and Sun, 1995; Salters and Stracke, 2004].
The recycled HIMU component with high H2O and CO2 is also inferred to have high F. While F and Cl are
likely elevated in the subducting slab, most of the subduction input is expected to be stripped off during
prograde dehydration, as reflected in elevated F and Cl concentrations in many arc magmas and forearc serpentines [e.g., Straub and Layne, 2003; John et al., 2011; Rose-Koga et al., 2012; Kendrick et al., 2013b; Debret
et al., 2014]. The data from Mangaia melt inclusions suggest that the recycled component in the HIMU
source is characterized by the fractionation of F from Cl, as Mangaia samples have Cl/Nb ratios similar to
ambient mantle, while F/Nd and H2O/Ce ratios remain higher than ambient mantle. However, the precise
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4462
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
mechanism responsible for fractionation of F and H2O from Cl in subduction zones is not well known. The
hydroxyl site in silicate minerals—which has a strong affinity for F [e.g., Smith, 1981; Smith et al., 1981; Wu
and Koga, 2013]—provides an ideal mineralogical site to promote recycling of H2O and F, but not Cl. Thus,
the residual mineralogy of the subducted slab after it has gone through prograde dehydration/melting,
including phlogopite, humite, apatite, and/or amphiboles (together with their high pressure polymorphs),
will play an important role in the fractionation of H2O and F from Cl [Rose-Koga et al., 2014].
5. Conclusions
Pb-isotopic studies of lavas from Mangaia have established this locality to be the surface expression of the
HIMU mantle end-member, which has long been considered to contain subducted oceanic crust in the
mantle source [Hofmann and White, 1982; Zindler and Hart, 1986; Hauri and Hart, 1993; Kelley et al., 2005;
Hanyu et al., 2011]. Sulfur isotopic work on melt inclusions from the island has established that the material
erupted at Mangaia was once at the Earth’s surface > 2.45 Ga, confirming the recycled origin hypothesis for
the HIMU mantle [Cabral et al., 2013].
Previous work on HIMU melt inclusions has been restricted to the study of crystalline inclusions from Mangaia
[Saal et al., 1998; Yurimoto et al., 2004; Paul et al., 2011], which cannot be used to measure volatile abundances. Additionally, because of the lack of dredging expeditions to this location, the scientific community has
been without glasses from Mangaia on which to perform volatile measurements. Here, we have homogenized
crystalline melt inclusions hosted in olivine phenocrysts from the island. We have performed in situ volatile,
trace, and major element analyses in the melt inclusions, in addition to in situ Pb-isotopic analyses. We find
Mangaia melt inclusions host elevated H2O (up to 1.73 wt %) and CO2 (up to 2346 ppm) contents.
We suggest that the H2O abundance in the HIMU mantle source is 440 ppm, which can generate the H2O
abundance observed in the melt inclusion with the highest H2O concentration when the melt fraction is
3%. The inclusions have relatively high concentrations of CO2 (up to 2346 ppm), which is consistent with
previous observations of carbonatite in Mangaia melt inclusions. The high levels of CO2 in HIMU lavas, and
hence the HIMU mantle source, are consistent with the experimental studies that require a carbonated protolith to generate the unique major element composition of the melts.
We also find the F/Nd ratios of the melt inclusions (30 6 9; 2r standard deviation) to be higher than the
canonical value (21), while Cl/Nb ratios in Mangaia inclusions are likely similar to the MORB mantle. We suggest that this may owe to the influence of hydroxyl sites in silicate minerals, which may promote recycling
of H2O and F, but not Cl.
Acknowledgments
M.J. acknowledges NSF grants EAR1145202, EAR-1348082, EAR-1347377,
and OCE-1153894 that supported this
work. E.F.R.-K. thanks the European
Synthesys FP7 ‘‘Capacities’’ Specific
Program for financing part of the
analytical cost of this research.
K.T.K. acknowledges French ANR grant
ANR-09-BLAN-038 (project SlabFlux)
that supported this work. The Nordsim
facility is funded and operated as a
joint Nordic research infrastructure
under an agreement with NOS-N. This
is Laboratory of Excellence ClerVolc
(ANR-10-LABX-006) contribution 126
and Nordsim laboratory contribution
376. We thank J.E. Dixon, P.J. le Roux,
A.E. Saal, and P. Cartigny for advice in
compiling the OIB and MORB database
that was used for creating many of the
figures. We also thank P.S. Hall for his
assistance in clarifying the diffusion
model. We acknowledge constructive
reviews from Bill White, John Lassiter,
and an anonymous reviewer.
CABRAL ET AL.
References
Allègre, C. J. (1982), Chemical geodynamics, Tectonophysics, 81, 109–132.
Alt, J. C., and D. A. H. Teagle (1999), The uptake of carbon during alteration of ocean crust, Geochim. Cosmochim. Acta, 63, 1527–1535.
Ardia, P., M. M. Hirschmann, A. C. Withers, and T. J. Tenner (2012), H2O storage capacity of olivine at 5–8GPa and consequences for dehydration partial melting of the upper mantle, Earth Planet. Sci. Lett., 345–348, 104–116, doi:10.1016/j.epsl.2012.05.038.
Asimow, P. D., and C. H. Langmuir (2003), The importance of water to oceanic mantle melting regimes, Nature, 421, 815–820.
Bach, W., J. C. Alt, Y. Niu, S. E. Humphris, J. Erzinger, and H. J. B. Dick (2001), The geochemical consequences of late-stage low-grade alteration of lower ocean crust at the SW Indian Ridge: Results from ODP Hole 735B (Leg 176), Geochim. Cosmochim. Acta, 65, 3267–3287.
Bucholz, C. E., G. A. Gaetani, M. D. Behn, and N. Shimizu (2013), Post-entrapment modification of volatiles and oxygen fugacity in olivinehosted melt inclusions, Earth Planet. Sci. Lett., 374, 145–155, doi:10.1016/j.epsl.2013.05.033.
Cabral, R. A., M. G. Jackson, E. F. Rose-Koga, K. T. Koga, M. J. Whitehouse, M. A. Antonelli, J. Farquhar, J. M. D. Day, and E. H. Hauri (2013),
Anomalous sulphur isotopes in plume lavas reveal deep mantle storage of Archaean crust, Nature, 496, 490–493, doi:10.1038/
nature12020.
Cartigny, P., F. Pineau, C. Aubaud, and M. Javoy (2008), Towards a consistent mantle carbon flux estimate: Insights from volatile systematics
(H2O/Ce, dD, CO2/Nb) in the North Atlantic mantle (14 N and 34 N), Earth Planet. Sci. Lett., 265, 672–685, doi:10.1016/
j.epsl.2007.11.011.
Castillo, P. R. (2014), Recycling of marine carbonate as a possible solution to the Pb paradox, Mineral. Mag., 358.
Chan, L.-H., J. C. Lassiter, E. H. Hauri, S. R. Hart, and J. Blusztajn (2009), Lithium isotope systematics of lavas from the Cook-Austral Islands:
Constraints on the origin of HIMU mantle, Earth Planet. Sci. Lett., 277, 433–442, doi:10.1016/j.epsl.2008.11.009.
Chase, C. G. (1981), Oceanic island Pb: Two-stage histories and mantle evolution, Earth Planet. Sci. Lett., 52, 277–284.
Chauvel, C., A. Hofmann, and P. Vidal (1992), HIMU-EM: The French Polynesian connection, Earth Planet. Sci. Lett., 110, 99–119.
Chauvel, C., W. McDonough, G. Guille, R. Maury, and R. Duncan (1997), Contrasting old and young volcanism in Rurutu Island, Austral chain,
Chem. Geol., 139, 125–143.
Chen, Y., A. Provost, P. Schiano, and N. Cluzel (2011), The rate of water loss from olivine-hosted melt inclusions, Contrib. Mineral. Petrol.,
162(3), 625–636, doi:10.1007/s00410-011-0616-5.
C 2014. American Geophysical Union. All Rights Reserved.
V
4463
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Cherniak, D. J. (2010), Diffusion in accessory minerals: Zircon, titanite, apatite, monazite and xenotime, Rev. Mineral. Geochem., 72(1), 827–
869, doi:10.2138/rmg.2010.72.18.
Clague, D. A., J. G. Moore, J. E. Dixon, and W. B. Friesen (1995), Petrology of submarine lavas from Kilauea’s Puna Ridge, Hawaii, J. Petrol., 36,
299–349.
Cohen, R. S., and R. K. O’Nions (1982), Identification of recycled continental material in the mantle from Sr, Nd and Pb isotope investigations, Earth Planet. Sci. Lett., 61, 73–84.
Dalou, C. (2011), Fluorine and chlorine fractionation in the sub-arc mantle: An experimental investigation, thesis, Univ. Blaise Pascal,
Clermont-Ferrand, France.
Danyushevsky, L. V., A. V. Sobolev, and N. N. Kononkova (1992), Methods of studying melt inclusions in minerals during investigations on
water-bearing primitive mantle melts (Tonga Trench boninites), Geochem. Int., 29, 48–62.
Danyushevsky, L. V., F. N. Della-Pasqua, and S. Sokolov (2000), Re-equilibration of melt inclusions trapped by magnesian olivine phenocrysts from subduction-related magmas: Petrological implications, Contrib. Mineral. Petrol., 138, 68–83.
Dasgupta, R., M. M. Hirschmann, and A. C. Withers (2004), Deep global cycling of carbon constrained by the solidus of anhydrous, carbonated eclogite under upper mantle conditions, Earth Planet. Sci. Lett., 227, 73–85, doi:10.1016/j.epsl.2004.08.004.
Dasgupta, R., M. M. Hirschmann, and N. D. Smith (2007), Partial melting experiments of peridotite 1 CO2 at 3 GPa and genesis of Alkalic
Ocean Island Basalts, J. Petrol., 48(11), 2093–2124, doi:10.1093/petrology/egm053.
Dasgupta, R., M. M. Hirschmann, W. F. McDonough, M. Spiegelman, and A. C. Withers (2009), Trace element partitioning between garnet
lherzolite and carbonatite at 6.6 and 8.6 GPa with applications to the geochemistry of the mantle and of mantle-derived melts, Chem.
Geol., 262, 57–77, doi:10.1016/j.chemgeo.2009.02.004.
Day, J. M. D., D. G. Pearson, C. G. Macpherson, D. Lowry, and J.-C. Carracedo (2009), Pyroxenite-rich mantle formed by recycled oceanic
lithosphere: Oxygen-osmium isotope evidence from Canary Island lavas, Geology, 37(6), 555–558, doi:10.1130/G25613A.1.
Day, J. M. D., D. G. Pearson, C. G. Macpherson, D. Lowry, and J. C. Carracedo (2010), Evidence for distinct proportions of subducted oceanic
crust and lithosphere in HIMU-type mantle beneath El Hierro and La Palma, Canary Islands, Geochim. Cosmochim. Acta, 74(22), 6565–
6589, doi:10.1016/j.gca.2010.08.021.
Debret, B., K. T. Koga, C. Nicollet, M. Andreani, and S. Schwartz (2014), F, Cl and S input via serpentinite in subduction zones: Implications
for the nature of the fluid released at depth, Terra Nova, 26(2), 96–101, doi:10.1111/ter.12074.
Devey, C. W., F. Albarede, J.-L. Chemin
ee, A. Michard, R. M€
uhe, and P. Stoffers (1990), Active submarine volcanism on the society hotspot
swell (west Pacific): A geochemical study, J. Geophys. Res., 95, 5049–5066.
Dixon, J., D. A. Clague, P. Wallace, and R. Poreda (1997), Volatiles in alkalic basalts form the North Arch Volcanic Field, Hawaii: Extensive
degassing of deep submarine-erupted alkalic series lavas, J. Petrol., 38(7), 911–939.
Dixon, J. E. (1997), Degassing of alkalic basalts, Am. Mineral., 82, 368–378.
Dixon, J. E., and D. A. Clague (2001), Volatiles in basaltic glasses from Loihi Seamount, Hawaii: Evidence for a relatively dry plume component, J. Petrol., 42, 627–654.
Dixon, J. E., L. Leist, C. Langmuir, and J.-G. Schilling (2002), Recycled dehydrated lithosphere observed in plume-influenced mid-oceanridge basalt, Nature, 420, 385–389.
Dosso, L., B. B. Hanan, H. Bougault, J.-G. Schilling, and J.-L Joron (1991), Sr-Nd-Pb geochemical morphology between 10 and 17 N on the
Mid-Atlantic Ridge: A new MORB isotope signature, Earth Planet. Sci. Lett., 106, 29–43.
Dosso, L., H. Bougault, C. Langmuir, C. Bollinger, O. Bonnier, and J. Etoubleau (1999), The age and distribution of mantle heterogeneity
along the Mid-Atlantic Ridge (31–41 N), Earth Planet. Sci. Lett., 170, 269–286.
Douglass, J., J.-G. Schilling, and R. H. Kingsley (1995), Influence of the discovery and Shona mantle plumes on the southern Mid-Atlantic
Ridge: Rare earth evidence, Geophys. Res. Lett., 22, 2893–2896.
Douglass, J., J.-G. Schilling, and D. Fontignie (1999), Plume-ridge interactions of the discovery and Shona mantle plumes with the southern
Mid-Atlantic Ridge (40 –55 S), J. Geophys. Res., 104, 2941–2962.
Dupuy, C., H. G. Barsczus, J. Dostal, P. Vidal, and J.-M. Liotard (1989), Subducted and recycled lithosphere as the mantle source of ocean
island basalts from southern Polynesia, central Pacific, Chem. Geol., 77, 1–18.
Eggler, D. (1978), The effect of CO2 upon partial melting of peridotite in the system Na2O-CaO-Al2O3-MgO-SiO2-CO2 to 35 Kb, with an analysis of melting in peridotite-H2O-CO2 system, Am. J. Sci., 278, 305–343.
Eiler, J. M., K. A. Farley, J. W. Valley, E. M. Stolper, E. H. Hauri, and H. Craig (1995), Oxygen isotope evidence against bulk recycled sediment
in the mantle sources of Pitcairn Island lavas, Nature, 377, 138–141.
Eisele, J., M. Sharma, S. J. G. Galer, J. Blichert-Toft, C. W. Devey, and A. W. Hofmann (2002), The role of sediment recycling in EM-1 inferred
from Os, Pb, Hf, Nd, Sr isotope and trace element systematics of the Pitcairn hotspot, Earth Planet. Sci. Lett., 196, 197–212.
Fontignie, D., and J.-G. Schilling (1991), 87Sr/86Sr and REE variations along the Easter Microplate boundaries (south Pacific): Application of
multivariate statistical analyses to ridge segmentation, Chem. Geol., 89, 209–241.
Frey, F. A., D. Clague, J. J. Mahoney, and J. M. Sinton (2000), Volcanism at the edge of the Hawaiian plume: Petrogenesis of submarine
alkalic lavas from the North Arch volcanic field, J. Petrol., 41, 667–691.
Gaetani, G. A., and T. L. Grove (1998), The influence of water on melting of mantle peridotite, Contrib. Mineral. Petrol., 131, 323–346.
Gaetani, G. A., and E. B. Watson (2000), Open system behavior of olivine-hosted melt inclusions, Earth Planet. Sci. Lett., 183, 27–41.
Gaetani, G. A., J. A. O’Leary, N. Shimizu, C. E. Bucholz, and M. Newville (2012), Rapid reequilibration of H2O and oxygen fugacity in olivinehosted melt inclusions, Geology, 40(10), 915–918, doi:10.1130/G32992.1.
Gale, A., M. Laubier, S. Escrig, and C. H. Langmuir (2013), Constraints on melting processes and plume-ridge interaction from comprehensive study of the FAMOUS and North Famous segments, Mid-Atlantic Ridge, Earth Planet. Sci. Lett., 365, 209–220, doi:10.1016/
j.epsl.2013.01.022.
Gast, P. W., G. R. Tilton, and C. Hedge (1964), Isotopic composition of lead and strontium from Ascension and Gough Islands, Science, 145,
1181–1185.
Gerbode, C., and R. Dasgupta (2010), Carbonate-fluxed melting of MORB-like pyroxenite at 2.9 GPa and genesis of HIMU Ocean Island
Basalts, J. Petrol., 51(10), 2067–2088, doi:10.1093/petrology/egq049.
Gillis, K. M., and L. A. Coogan (2011), Secular variation in carbon uptake into the ocean crust, Earth Planet. Sci. Lett., 302, 385–392, doi:
10.1016/j.epsl.2010.12.030.
Graham, D., S. Humphris, and W. Jenkins (1992), Helium isotope geochemistry of some volcanic rocks from Saint Helena, Earth Planet. Sci.
Lett., 110, 121–131.
Hacker, B. R. (2008), H2O subduction beyond arcs, Geochem. Geophys. Geosyst., 9, Q03001, doi:10.1029/2007GC001707.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4464
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Hanyu, T., and I. Kaneoka (1997), The uniform and low 3He/4He ratios of HIMU basalts as evidence for their origin as recycled materials,
Nature, 390, 273–276.
Hanyu, T., et al. (2011), Geochemical characteristics and origin of the HIMU reservoir: A possible mantle plume source in the lower mantle,
Geochem. Geophys. Geosyst., 12, Q0AC09, doi:10.1029/2010GC003252.
Hanyu, T., et al. (2014), Isotope evolution in the HIMU reservoir beneath St. Helena: Implications for the mantle recycling of U and Th, Geochim. Cosmochim. Acta, 143, 232–252, doi:10.1016/j.gca.2014.03.016.
Hart, S. (2011), The Mantle Zoo: New species, endangered species, extinct species, Mineral. Mag., 3, 983.
Hart, S., and H. Staudigel (1982), The control of alkalies and uranium in seawater by ocean crust alteration, Earth Planet. Sci. Lett., 58, 202–
212.
Hart, S. R. (1988), Heterogeneous mantle domains: Signatures, genesis and mixing chronologies, Earth Planet. Sci. Lett., 90, 273–296.
Hart, S. R., E. H. Hauri, L. A. Oschmann, and J. A. Whitehead (1992), Mantle plumes and entrainment: Isotopic evidence, Science, 256, 517–
520.
Hauri, E. (2002), SIMS analysis of volatiles in silicate glasses. 2: Isotopes and abundances in Hawaiian melt inclusions, Chem. Geol., 183, 115–
141.
Hauri, E. H., and S. R. Hart (1993), Re-Os isotope systematics of HIMU and EMII oceanic island basalts from the south Pacific Ocean, Earth
Planet. Sci. Lett., 114, 353–371.
Hauri, E. H., and S. R. Hart (1997), Rhenium abundances and systematics in oceanic basalts, Chem. Geol., 139, 185–205.
Hauri, E., N. Shimizu, J. Dieu, and S. Hart (1993), Evidence for hotspot-related carbonatite metasomatism in the oceanic upper mantle,
Nature, 365, 221–227.
H
emond, C., C. Devey, and C. Chauvel (1994), Source compositions and melting processes in the Society and Austral plumes (South Pacific
Ocean): Element and isotope (Sr, Nd, Pb, Th) geochemistry, Chem. Geol., 115, 7–45.
Herzberg, C. (2011), Identification of source lithology in the Hawaiian and Canary Islands: Implications for origins, J. Petrol., 52(1), 113–146,
doi:10.1093/petrology/egq075.
Herzberg, C., and E. Gazel (2009), Petrological evidence for secular cooling in mantle plumes, Nature, 458, 619–622, doi:10.1038/
nature07857.
Herzberg, C., R. A. Cabral, M. G. Jackson, C. Vidito, J. M. D. Day, and E. H. Hauri (2014), Phantom Archean crust in Mangaia hotspot lavas and
the meaning of heterogeneous mantle, Earth Planet. Sci. Lett., 396, 97–106.
Hirschmann, M. M. (2006), Water, melting, and the deep Earth H2O cycle, Annu. Rev. Earth Planet. Sci., 34, 629–653, doi:10.1146/10.1146/
annurev.earth.34.031405.125211.
Hirschmann, M. M., T. Tenner, C. Aubaud, and A. C. Withers (2009), Dehydration melting of nominally anhydrous mantle: The primacy of
partitioning, Phys. Earth Planet. Inter., 176, 54–68, doi:10.1016/j.pepi.2009.04.001.
Hirth, G., and D. L. Kohlstedt (1996), Water in the oceanic upper mantle: Implications for rheology, melt extraction and the evolution of the
lithosphere, Earth Planet. Sci. Lett., 144, 93–108.
Hoernle, K., G. Tilton, M. J. Le Bas, S. Duggen, and D. Garbe-Sch€
onberg (2002), Geochemistry of oceanic carbonatites compared with continental carbonatites: Mantle recycling of oceanic crustal carbonate, Contrib. Mineral. Petrol., 142, 520–542.
Hofmann, A. W. (1997), Mantle geochemistry: The message from oceanic volcanism, Nature, 385, 219–229.
Hofmann, A. W., and S. R. Hart (1978), An assessment of local and regional isotopic equilibrium in the mantle, Earth Planet. Sci. Lett., 38,
44–62.
Hofmann, A. W., and W. M. White (1982), Mantle plumes from ancient oceanic crust, Earth Planet. Sci. Lett., 57, 421–436.
Honda, M., and J. D. Woodhead (2005), A primordial solar-neon enriched component in the source of EM-I-type ocean island basalts from
the Pitcairn Seamounts, Polynesia, Earth Planet. Sci. Lett., 236, 597–612, doi:10.1016/j.epsl.2005.05.038.
Jackson, M. G., and R. Dasgupta (2008), Compositions of HIMU, EM1, and EM2 from global trends between radiogenic isotopes and major
elements in ocean island basalts, Earth Planet. Sci. Lett., 276, 175–186, doi:10.1016/j.epsl.2008.09.023.
Jackson, M. G., and A. M. Jellinek (2013), Major and trace element composition of the high 3He/4He mantle: Implications for the composition of a nonchonditic Earth, Geochem. Geophys. Geosyst., 14, 2954–2976, doi:10.1002/ggge.20188.
Jackson, M. G., S. R. Hart, A. A. P. Koppers, H. Staudigel, J. Konter, J. Blusztajn, M. Kurz, and J. A. Russell (2007), The return of subducted continental crust in Samoan lavas, Nature, 448, 684–687, doi:10.1038/nature06048.
Jackson, M. G., R. W. Carlson, M. D. Kurz, P. D. Kempton, D. Francis, and J. Blusztajn (2010), Evidence for the survival of the oldest terrestrial
mantle reservoir, Nature, 466, 853–856, doi:10.1038/nature09287.
Jackson, M. G., D. Weis, and S. Huang (2012), Major element variations in Hawaiian shield lavas: Source features and perspectives from
global ocean island basalt (OIB) systematics, Geochem. Geophys. Geosyst., 13, Q09009, doi:10.1029/2012GC004268.
Jambon, A., B. D
eruelle, G. Dreibus, and F. Pineau (1995), Chlorine and bromine abundance in MORB: The contrasting behaviour of the
Mid-Atlantic Ridge and East Pacific Rise and implications for chlorine geodynamic cycle, Chem. Geol., 126, 101–117.
John, T., G. D. Layne, K. M. Haase, and J. D. Barnes (2010), Chlorine isotope evidence for crustal recycling into the Earth’s mantle, Earth
Planet. Sci. Lett., 298(1–2), 175–182, doi:10.1016/j.epsl.2010.07.039.
John, T., M. Scambelluri, M. Frische, J. D. Barnes, and W. Bach (2011), Dehydration of subducting serpentinite: Implications for halogen
mobility in subduction zones and the deep halogen cycle, Earth Planet. Sci. Lett., 308(1–2), 65–76, doi:10.1016/j.epsl.2011.05.038.
Kawabata, H., T. Hanyu, Q. Chang, J.-I. Kimura, A. R. L. Nichols, and Y. Tatsumi (2011), The Petrology and geochemistry of St. Helena alkali
basalts: Evaluation of the oceanic crust-recycling model for HIMU OIB, J. Petrol., 52(4), 791–838, doi:10.1093/petrology/egr003.
Kelemen, P. B., G. M. Yogodzinski, and D. W. Scholl (2003), Along-strike variation in lavas of the Aleutian island arc: Implications for the genesis of high Mg# andesite and the continental crust. Inside the Subduction Factory, Geophys. Monogr. Ser, 138, 223–276.
Kelley, K. A., T. Plank, J. Ludden, and H. Staudigel (2003), Composition of altered oceanic crust at ODP Sites 801 and 1149, Geochem. Geophys. Geosyst., 4(6), 8910, doi:10.1029/2002GC000435.
Kelley, K. A., T. Plank, L. Farr, J. Ludden, and H. Staudigel (2005), Subduction cycling of U, Th, and Pb, Earth Planet. Sci. Lett., 234, 369–383,
doi:10.1016/j.epsl.2005.03.005.
Kelley, K. A., R. Kingsley, and J.-G. Schilling (2013), Composition of plume-influenced mid-ocean ridge lavas and glasses from the MidAtlantic Ridge, East Pacific Rise, Galapagos Spreading Center, and Gulf of Aden, Geochem. Geophys. Geosyst., 14, 223–242, doi:10.1029/
2012GC004415.
Kendrick, M. A., R. Arculus, P. Burnard, and M. Honda (2013a), Quantifying brine assimilation by submarine magmas: Examples from the
Galapagos Spreading Centre and Lau Basin, Geochim. Cosmochim. Acta, 123, 150–165, doi:10.1016/j.gca.2013.09.012.
Kendrick, M. A., M. Honda, T. Pettke, M. Scambelluri, D. Phillips, and A. Giuliani (2013b), Subduction zone fluxes of halogens and noble
gases in seafloor and forearc serpentinites, Earth Planet. Sci. Lett., 365, 86–96, doi:10.1016/j.epsl.2013.01.006.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4465
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Kendrick, M. A., M. G. Jackson, A. J. R. Kent, E. H. Hauri, P. J. Wallace, and J. Woodhead (2014), Contrasting behaviours of CO2, S, H2O and
halogens (F, Cl, Br, and I) in enriched-mantle melts from Pitcairn and Society seamounts, Chem. Geol., 370, 69–81, doi:10.1016/
j.chemgeo.2014.01.019.
Kent, A. J., M. D. Norman, I. D. Hutcheon, and E. M. Stolper (1999a), Assimilation of seawater-derived components in an oceanic volcano:
Evidence from matrix glasses and glass inclusions from Loihi seamount, Hawaii, Chem. Geol., 156, 299–319.
Kent, A. J., D. A. Clague, M. Honda, E. M. Stolper, I. D. Hutcheon, and M. D. Norman (1999b), Widespread assimilation of a seawater-derived
component at Loihi Seamount, Hawaii, Geochim. Cosmochim. Acta, 63, 2749–2761.
Kingsley, R. H. (2002), The geochemistry of basalts from the Easter Microplate boundaries and the Salas y Gomez seamount chain: A comprehensive study of mantle plume-spreading center interaction, dissertation, Univ. of Rhode Island, Narragansett, R. I.
Kingsley, R. H., and J.-G. Schilling (1998), Plume-ridge interaction in the Easter-Salas y Gomez seamount chain-Easter Microplate system: Pb
isotope evidence, J. Geophys. Res., 103, 24,159–24,177.
Kogiso, T., Y. Tatsumi, G. Shimoda, and H. G. Barsczus (1997), High l (HIMU) ocean island basalts in southern Polynesia: New evidence for
whole mantle scale recycling of subducted oceanic crust, J. Geophys. Res., 102, 8085–8103.
Koleszar, A. M., A. E. Saal, E. H. Hauri, A. N. Nagle, Y. Liang, and M. D. Kurz (2009), The volatile contents of the Galapagos plume: Evidence
for H2O and F open system behavior in melt inclusion, Earth Planet. Sci. Lett., 287, 442–452, doi:10.1016/j.epsl.2009.08.029.
Krienitz, M.-S., C.-D. Garbe-Sch€
onberg, R. L. Romer, A. Meixner, K. M. Haase, and N. A. Stroncik (2012), Lithium isotope variations in Ocean
Island Basalts—Implications for the development of mantle heterogeneity, J. Petrol., 53(11), 2333–2347, doi:10.1093/petrology/egs052.
Kushiro, I. (1975), On the nature of silicate melt and its significance in magma genesis: Regularities in the shift of the liquidus boundaries
involving olivine, pyroxene, and silica minerals, Am. J. Sci., 275, 411–431.
Langmuir, C., et al. (1997), Hydrothermal vents near a mantle hot spot: The Lucky Strike vent field at 37 N on the Mid-Atlantic Ridge, Earth
Planet. Sci. Lett., 148, 69–91.
Lassiter, J. C., E. H. Hauri, I. K. Nikogosian, and H. G. Barsczus (2002), Chlorine–potassium variations in melt inclusions from Raivavae and
Rapa, Austral Islands: Constraints on chlorine recycling in the mantle and evidence for brine-induced melting of oceanic crust, Earth
Planet. Sci. Lett., 202, 525–540.
Lassiter, J. C., J. Blichert-Toft, E. H. Hauri, and H. G. Barsczus (2003), Isotope and trace element variations in lavas from Raivavae and Rapa,
Cook-Austral Islands: Constraints on the nature of HIMU- and EM-mantle and the origin of mid-plate volcanism in French Polynesia,
Chem. Geol., 202, 115–138, doi:10.1016/j.chemgeo.2003.08.002.
le Roux, P. J., A. P. le Roex, and J.-G. Schilling (2002), Crystallization processes beneath the southern Mid-Atlantic Ridge (40–55 S), evidence
for high-pressure initiation of crystallization, Contrib. Mineral. Petrol., 142, 582–602.
MacDonald, G. A., and T. Katsura (1965), Eruption of Lassen Peak, Cascade Range, California, in 1915: Example of mixed magmas, Geol. Soc.
Am. Bull., 76, 475–482.
Mackwell, S. J., and D. L. Kohlstedt (1990), Diffusion of hydrogen in olivine: Implications for water in the mantle, J. Geophys. Res., 95, 5079–
5088.
Mallik, A., and R. Dasgupta (2012), Reaction between MORB-eclogite derived melts and fertile peridotite and generation of ocean island
basalts, Earth Planet. Sci. Lett., 329–330, 97–108, doi:10.1016/j.epsl.2012.02.007.
McDonough, W. F., and S. S. Sun (1995), The composition of the Earth, Chem. Geol., 120, 223–253.
Metrich, N., V. Zanon, L. Creon, A. Hildenbrand, M. Moreira, and F. O. Marques (2014), Is the ‘‘Azores Hotspot’’ a Wetspot? Insights from the
Geochemistry of Fluid and Melt Inclusions in Olivine of Pico Basalts, J. Petrol, 55(2), 377–393, doi:10.1093/petrology/egt071.
Michael, P. (1995), Regionally distinctive sources of depleted MORB: Evidence from trace elements and H2O, Earth Planet. Sci. Lett., 131,
301–320.
Michael, P. J., and W. C. Cornell (1998), Influence of spreading rate and magma supply on crystallization and assimilation beneath midocean ridges: Evidence from chlorine and major element chemistry of mid-ocean ridge basalts, J. Geophys. Res., 103, 18,325–18,356.
Michael, P. J., and J.-G. Schilling (1989), Chlorine in mid-ocean ridge magmas: Evidence for assimilation of seawater-influenced components, Geochim. Cosmochim. Acta, 53, 3131–3143.
Nakamura, Y., and M. Tatsumoto (1988), Pb, Nd, and Sr isotopic evidence for a multicomponent source for rocks of Cook-Austral Islands
and heterogeneities of mantle plumes, Geochim. Cosmochim. Acta, 52, 2909–2924.
Nishio, Y., S. Nakai, T. Kogiso, and H. Barsczus (2005), Lithium, strontium, and neodymium isotopic compositions of oceanic island basalts
in the Polynesian region: Constraints on a Polynesian HIMU origin, Geochem. J., 39, 91–103.
O’Leary, J. A., G. A. Gaetani, and E. H. Hauri (2010), The effect of tetrahedral Al31 on the partitioning of water between clinopyroxene and
silicate melt, Earth Planet. Sci. Lett., 297(1–2), 111–120, doi:10.1016/j.epsl.2010.06.011.
Pan, Y., and R. Batiza (1998), Major element chemistry of volcanic glasses from the Easter Seamount Chain: Constraints on melting conditions in the plume channel, J. Geophys. Res., 103, 5287–5304.
Parai, R., S. Mukhopadhyay, and J. C. Lassiter (2009), New constraints on the HIMU mantle from neon and helium isotopic compositions of
basalts from the Cook-Austral Islands, Earth Planet. Sci. Lett., 277(1–2), 253–261, doi:10.1016/j.epsl.2008.10.014.
Paul, B., J. D. Woodhead, J. Hergt, L. Danyushevsky, T. Kunihiro, and E. Nakamura (2011), Melt inclusion Pb-isotope analysis by LA-MCICPMS: Assessment of analytical performance and application to OIB genesis, Chem. Geol., 289, 210–223, doi:10.1016/
j.chemgeo.2011.08.005.
Portnyagin, M., K. Hoernle, P. Plechov, N. Mironov, and S. Khubunaya (2007), Constraints on mantle melting and composition and nature of
slab components in volcanic arcs from volatiles (H2O, S, Cl, F) and trace elements in melt inclusions from the Kamchatka Arc, Earth
Planet. Sci. Lett., 255, 53–69, doi:10.1016/j.epsl.2006.12.005.
Rose-Koga, E. F., K. T. Koga, P. Schiano, M. Le Voyer, N. Shimizu, M. J. Whitehouse, and R. Clocchiatti (2012), Mantle source heterogeneity
for South Tyrrhenian magmas revealed by Pb isotopes and halogen contents of olivine-hosted melt inclusions, Chem. Geol., 334, 266–
279, doi:10.1016/j.chemgeo.2012.10.033.
Rose-Koga, E. F., K. T. Koga, M. Hamada, T. H
elouis, M. J. Whitehouse, and N. Shimizu (2014), Volatile (F and Cl) concentrations in Iwate
olivine-hosted melt inclusions indicating low-temperature subduction, Earth Planets Space, 66, 81.
Roy-Barman, M., and C. J. Allègre (1995), 187Os/186Os in oceanic island basalts: Tracing oceanic crust recycling in the mantle, Earth Planet.
Sci. Lett., 129, 145–161.
Saal, A. E., S. R. Hart, N. Shimizu, E. H. Hauri, and G. D. Layne (1998), Pb isotopic variability in melt inclusions from Oceanic Island Basalts,
Polynesia, Science, 282, 1481–1484, doi:10.1126/science.282.5393.1481.
Saal, A. E., E. H. Hauri, C. H. Langmuir, and M. R. Perfit (2002), Vapour undersaturation in primitive mid-ocean-ridge basalt and the volatile
content of Earth’s upper mantle, Nature, 419, 451–455.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4466
Geochemistry, Geophysics, Geosystems
10.1002/2014GC005473
Saal, A. E., S. R. Hart, N. Shimizu, E. H. Hauri, G. D. Layne, and J. M. Eiler (2005), Pb isotopic variability in melt inclusions from the EMI-EMIIHIMU mantle end-members and the role of the oceanic lithosphere, Earth Planet. Sci. Lett., 240, 605–620, doi:10.1016/j.epsl.2005.10.002.
Salters, V., and W., White (1998), Hf isotope constraints on mantle evolution, Chem. Geol., 145, 447–460.
Salters, V. J. M., and A. Sachi-Kocher (2010), An ancient metasomatic source for the Walvis Ridge basalts, Chem. Geol., 273, 151–167, doi:
10.1016/j.chemgeo.2010.02.010.
Salters, V. J. M., and A. Stracke (2004), Composition of the depleted mantle, Geochem. Geophys. Geosyst., 5, Q05004, doi:10.1029/
2003GC000597.
Salters, V. J. M., S. Mallick, S. R. Hart, C. E. Langmuir, and A. Stracke (2011), Domains of depleted mantle: New evidence from hafnium and
neodymium isotopes, Geochem. Geophys. Geosyst., 12, Q08001, doi:10.1029/2011GC003617.
Sarbas, B. (2008), The GEOROC database as part of a growing geoinformatics network, translated by S. R. Brady, A. K. Sinha, and L. C. Gundersen, Geoinformatics. Data to Knowledge, Proceedings, U. S. Geol. Surv. Sci. Invest. Rep., pp. 42–43.
Schiano, P., K. W. Burton, B. Dupr
e, J.-L. Birck, G. Guille, and C. J. Allègre (2001), Correlated Os-Pb-Nd-Sr isotopes in the Austral-Cook chain
basalts: The nature of mantle components in plume sources, Earth Planet. Sci. Lett., 186, 527–537.
Schilling, J.-G., H. Sigurdsson, A. N. Davis, and R. N. Hey (1985), Easter microplate evolution, Nature, 317, 325–331.
Shaw, A. M., M. D. Behn, S. E. Humphris, R. A. Sohn, and P. M. Gregg (2010), Deep pooling of low degree melts and volatile fluxes at the
85 E segment of the Gakkel Ridge: Evidence from olivine-hosted melt inclusions and glasses, Earth Planet. Sci. Lett., 289(3–4), 311–322,
doi:10.1016/j.epsl.2009.11.018.
Simons, K., J. Dixon, J.-G. Schilling, R. Kingsley, and R. Poreda (2002), Volatiles in basaltic glasses from the Easter-Salas y Gomez Seamount
Chain and Easter Microplate: Implications for geochemical cycling of volatile elements, Geochem. Geophys. Geosyst., 3, 1–29, doi:
10.1029/2001GC000173.
Smith, J. V. (1981), Halogen and phosphorus storage in the Earth, Nature, 289, 762–765.
Smith, J. V., J. S. Delaney, R. L. Hervig, and J. B. Dawson (1981), Storage of F and Cl in the upper mantle: Geochemical implications, Lithos,
14, 133–147.
Sobolev, A. V., and L. V. Danyushevsky (1994), Petrology and geochemistry of boninites from the north termination of the Tonga Trench:
Constraints on the generation conditions of primary high-carboninite magmas, J. Petrol., 35, 1183–1211.
Spilliaert, N., N. M
etrich, and P. Allard (2006), S-Cl-F degassing pattern of water-rich alkali basalt: Modelling and relationship with eruption
styles on Mount Etna volcano, Earth Planet. Sci. Lett., 248, 772–786, doi:10.1016/j.epsl.2006.06.031.
Staudigel, H., A. Zindler, S. Hart, and T. Leslie (1984), The isotope systematics of a juvenile intraplate volcano: Pb, Nd, and Sr isotope ratios
of basalts from Loihi Seamount, Hawaii, Earth Planet. Sci. Lett., 69, 13–29.
Staudigel, H., T. Plank, B. White, and H. Schmincke (1996), Geochemical fluxes during seafloor alteration of the basaltic upper oceanic crust:
DSDP Sites 417 and 418: Subduction: Top to Bottom, Geophys. Monogr., 96, 19–38.
Stracke, A., M. Bizimis, and V. J. M. Salters (2003), Recycling oceanic crust: Quantitative constraints, Geochem. Geophys. Geosyst., 4(3), 8003,
doi:10.1029/2001GC000223.
Straub, S. M., and G. D. Layne (2003), The systematics of chlorine, fluorine, and water in Izu arc front volcanic rocks: Implications for volatile
recycling in subduction zones, Geochim. Cosmochim. Acta, 67(21), 4179–4203, doi:10.1016/S0016–7037(03)00307-7.
Stroncik, N. A., and K. M. Haase (2004), Chlorine in oceanic intraplate basalts: Constraints on mantle sources and recycling processes, Geology, 32(11), 945, doi:10.1130/G21027.1.
Tatsumi, Y., K. Oguri, G. Shimoda, T. Kogiso, and H. G. Barsczus (2000), Contrasting behavior of noble-metal elements during magmatic differentiation in basalts from the Cook Islands, Polynesia, Geology, 28, 131–134.
van Keken, P. E., B. R. Hacker, E. M. Syracuse, and G. A. Abers (2011), Subduction factory. 4: Depth-dependent flux of H2O from subducting
slabs worldwide, J. Geophys. Res., 116, B01401, doi:10.1029/2010JB007922.
Van Orman, J. A., T. L. Grove, and N. Shimizu (2001), Rare earth element diffusion in diopside: Influence of temperature, pressure, and ionic
radius, and an elastic model for diffusion in silicates, Contrib. Mineral. Petrol., 141, 687–703, doi:10.1007/s004100100269.
Wallmann, K. (2001), The geological water cycle and the evolution of marine d18O values, Geochim. Cosmochim. Acta, 65, 2469–2485.
Weaver, B. L. (1991), The origin of ocean island basalt end-member compositions: Trace element and isotopic constraints, Earth Planet. Sci.
Lett., 104, 381–397.
White, W. M. (2010), Oceanic Island Basalts and mantle plumes: The geochemical perspective, Annu. Rev. Earth Planet. Sci., 38, 133–160, doi:
10.1146/annurev-earth-040809-152450.
White, W. M., and A. W. Hofmann (1982), Sr and Nd isotope geochemistry of oceanic basalts and mantle evolution, Nature, 296, 821–825.
Witham, F., J. Blundy, S. C. Kohn, P. Lesne, J. Dixon, S. V. Churakov, and R. Botcharnikov (2012), SolEx: A model for mixed COHSCl-volatile
solubilities and exsolved gas compositions in basalt, Comput. Geosci., 45, 87–97, doi:10.1016/j.cageo.2011.09.021.
Woodhead, J. D. (1996), Extreme HIMU in an oceanic setting: The geochemistry of Mangaia Island (Polynesia), and temporal evolution of
the Cook-Austral hotspot, J. Volcanol. Geotherm. Res., 72, 1–19.
Woodhead, J. D., and C. W. Devey (1993), Geochemistry of the Pitcairn seamounts. I: Source character and temporal trends, Earth Planet.
Sci. Lett., 116, 81–99.
Workman, R. K., S. R. Hart, M. Jackson, M. Regelous, K. A. Farley, J. Blusztajn, M. Kurz, and H. Staudigel (2004), Recycled metasomatized lithosphere as the origin of the Enriched Mantle II (EM2) end-member: Evidence from the Samoan Volcanic Chain, Geochem. Geophys. Geosyst., 5, Q04008, doi:10.1029/2003GC000623.
Workman, R. K., E. Hauri, S. R. Hart, J. Wang, and J. Blusztajn (2006), Volatile and trace elements in basaltic glasses from Samoa: Implications
for water distribution in the mantle, Earth Planet. Sci. Lett., 241, 932–951, doi:10.1016/j.epsl.2005.10.028.
Wu, J., and K. T. Koga (2013), Fluorine partitioning between hydrous minerals and aqueous fluid at 1GPa and 770–947 C: A new constraint
on slab flux, Geochim. Cosmochim. Acta, 119, 77–92, doi:10.1016/j.gca.2013.05.025.
Yang, H.-J., F. A. Frey, and D. A. Clague (2003), Constraints on the source components of lavas forming the Hawaiian North Arch and Honolulu Volcanics, J. Petrol., 44, 603–627.
Yurimoto, H., T. Kogiso, K. Abe, H. G. Barsczus, A. Utsunomiya, and S. Maruyama (2004), Lead isotopic compositions in olivine-hosted melt
inclusions from HIMU basalts and possible link to sulfide components, Phys. Earth Planet. Inter., 146, 231–242, doi:10.1016/
j.pepi.2003.08.013.
Zhang, Y. (2008), Geochemical Kinetics, pp. 173–324, Princeton Univ. Press, Princeton, N. J.
Zindler, A., and S. Hart (1986), Chemical geodynamics, Annu. Rev. Earth Planet. Sci., 14, 493–571.
Zindler, A., E. Jagoutz, and S. Goldstein (1982), Nd, Sr and Pb isotopic systematics in a three-component mantle: A new perspective, Nature,
298, 519–523.
CABRAL ET AL.
C 2014. American Geophysical Union. All Rights Reserved.
V
4467